Magnetic Stratigraphy (International Geophysics) (International Geophysics)

MAGNETIC STRATIGRAPHY

This is Volume 64 in the INTERNATIONAL GEOPHYSICS SERIES A series of monographs and textbooks Edited by RENATA DMOWSKA and JAMES R. HOLTON A complete list of books in this series appears at the end of this volume.

MAGNETIC STRATIGRAPHY NElL D. OPDYKE AND

JAMES E. T. CHANNELL

Department of Geology University of Florida Gainesville, Florida

ACADEMIC PRESS San Diego New York

Sydney

London

Boston

Tokyo

Toronto

This book is printed on acid-free paper. ( ~

Copyright 9 1996 by ACADEMIC PRESS All Rights Reserved. No part of this publication may be reproduced or transmitted in any form or by any means, electronic or mechanical, including photocopy, recording, or any information storage and retrieval system, without permission in writing from the publisher. A c a d e m i c P r e s s , Inc. 525 B Street, Suite 1900, San Diego, California 92101-4495, USA http ://w w w. apne t. com Academic Press Limited 24-28 Oval Road, London NWI 7DX, UK http://www.hbuk.co.uk/ap/ Library of Congress Cataloging-in-Publication Data Opdyke, N. D. Magnetic stratigraphy / by N.D. Opdyke, J.E.T. Channell. p. cm. Includes index. ISBN 0-12-527470-X (alk. paper) I. Channell, 1. Paleomagnetism. 2. Stratigraphic correlation. J. E.T. II. Title. QE501.4.P35063 1996 551.7'01--dc20 96-5925 CIP PRINTED IN THE UNITED STATES OF AMERICA 96 97 98 99 00 01 BC 9 8 7 6

5

4

3

2

1

In memory of S. K. Runcorn 1922-1995

a This Page Intentionally Left Blank

Contents

Preface Table Key

xi xiii

1

Introduction and History 1.1 Introduction 1.2 Early Developments 1.3 Evidence for Field Reversal

2

The Earth's Magnetic Field

2.1 2.2 2.3 2.4 2.5

Introduction The Dipole Hypothesis Models of Field Reversal Polarity Transition Records and VGP PalEs Statistical Structure of the Geomagnetic Polarity Pattern

9 15 17 20 24 o,

VII

eeo

VIII

Contents

3

Magnetization Processes and Magnetic Properties of Sediments 3.1 3.2 3.3 3.4 3.5

Basic Principle Magnetic Minerals Magnetization Processes Magnetic Properties of Marine Sediments Magnetic Properties of Terrestrial Sediments

26 29 34 37 45

4

Laboratory Techniques 4.1 4.2 4.3 4.4

Introduction Resolving Magnetization Components Statistics Practical Guide to the Identification of Magnetic Minerals

49 50 58 61

5

Fundamentals of Magnetic Stratigraphy 5.1 5.2 5.3 5.4 5.5 5.6 5.7

Principles and Definitions Polarity Zone and Polarity Chron Nomenclature Field Tests for Timing of Remanence Acquisition Field Sampling for Magnetic Polarity Stratigraphy Presentation of Magnetostratigraphic Data Correlation of Polarity Zones to the GPTS Quality Criteria for Magnetos~atigraphic Data

74 75 82 85 91 91 93

6

The Pliocene-Pleistocene Polarity Record

6.1 6.2 6.3 6.4 6.5

Early Development of the Plio-Pleistocene OPTS Subchrons within the Maluyama Chron Magnetic Stratigraphy in Plio-Pleistocene Sediments Astrochronologic Calibration of the Plio-Pleistocene GPTS 4~ Age Calibration of the Plio-Pleistocene GPTS

95 96 99 107 111

Contents

ix

7

Late Cretaceous-Cenozoic GPTS 7.1 Oceanic Magnetic Anomaly Record 7.2 Numerical Age Control

113 119

8

Paleogene and Miocene Marine Magnetic Stratigraphy 8.1 Miocene Magnetic Stratigraphy 8.2 Paleogene Magnetic Stratigraphy 8.3 Integration of Chemostratigraphy and Magnetic Stratigraphy

126 132 136

9

Cenozoic Terrestrial Magnetic Stratigraphy

9.1 9.2 9.3 9.4 9.5 9.6

Introduction North American Neogene and Quaternary Eurasian Neogene African and South American Neogene North American and Eurasian Paleogene Mammal Dispersal in the Northern Hemisphere

144 145 146 155 156 160

10

Jurassic-Early Cretaceous GPTS 10.1 10.2 10.3 10.4

Oceanic Magnetic Anomaly Record Numerical Age Control Oxfordian-Aptian Time Scales Hettangian-Oxfordian Time Scales

168 171 175 176

11 Jurassic and Cretaceous Magnetic Stratigraphy 11.1 CretaceousMagnetic Stratigraphy 11.2 Jurassic Magnetic Stratigraphy 11.3 Correlation of Late Jurassic-Cretaceous Stage Boundaries to the GPTS

182 196 202

x

Contents

12

Triassic and Paleozoic Magnetic Stratigraphy 12.1 12.2 12.3 12.4 12.5 12.6

Introduction Triassic Permian Carboniferous Pre-Carboniferous Polarity Bias in the Phanerozoic

205 207 215 220 222 229

13

Secular Variation and Brunhes Chron Excursions 13.1 Introduction 13.2 Sediment Records of Secular Variation 13.3 Geomagnetic Excursions in the Brunhes Chron

233 234 242

14

Rock Magnetic Stratigraphy and Paleointensities 14.1 Introduction 14.2 Magnetic Parameters Sensitive to Concentration, Grain Size, and Mineralogy 14.3 Rock Magnetic Stratigraphy in Marine Sediments 14.4 Rock Magnetic Stratigraphy in Loess Deposits 14.5 Rock Magnetic Stratigraphy in Lake Sediments 14.6 Future Prospects for Rock Magnetic Stratigraphy 14.7 Paleointensity Determinations

251 253 265 266 269 270

Bibliography

277

Index

333

250

Preface

Magnetic polarity stratigraphy, the stratigraphic record of polarity reversals in rocks and sediments, is now thoroughly integrated into biostratigraphy and chemostratigraphy. For Late Mesozoic to Quaternary times, the geomagnetic polarity record is central to the construction of geologic time scales, linking biostratigraphies, isotope stratigraphies, and absolute ages. The application of magnetic stratigraphy in geologic investigations is now commonplace; however, the use of magnetic stratigraphy as a correlation tool in sediments and lava flows has developed only in the past 35 years. As recently as the mid-1960s, magnetic polarity stratigraphy dealt with only the last 3 My of Earth history; today, a fairly well-resolved magnetic polarity sequence has been documented back to 300 My before present. An increasing number of scientists in allied fields now use and apply magnetic stratigraphy. This book is aimed at this expanding practitioner base, providing information about the principles of magnetostratigraphy and the present state of our knowledge concerning correlations among the various (biostratigraphic, chemostratigraphic, magnetostratigraphic, and numerical) facets of geologic time. Magnetic Stratigraphy documents the historical development of magnetostratigraphy (Chapter 1), the principal characteristics of the Earth's magnetic field (Chapter 2), the magnetization process and the magnetic properties of sediments (Chapter 3), and practical laboratory techniques (Chapter 4). Chapter 5 describes the essential elements that should be incorporated into any magnetostratigraphic investigation. The middle of the book is devoted to a survey of the status of magnetic stratigraphy in Jurassic through Quaternary marine (Chapters 6-8, 10, and 11) and terrestrial (Chapter 9) sedimentary rocks and describes the integration of the biostratigraphic and chemostratigraphic record with the geomagnetic polarity time scale (GPTS). Recent developments in extending the magnetic polarity record to pre-Jurassic sediments and the emergence xi

oo

XII

Preface

of a magnetic polarity record for the Triassic, Permian, and Carboniferous periods are chronicled in Chapter 12. The record of secular variation of the geomagnetic field has been important in attempting correlations in sediments that span the last 10,000 years (Chapter 13). Magnetic stratigraphy is expanding further through the use of nondirectional magnetic properties of sediments, such as magnetic susceptibility and a host of other magnetic parameters sensitive to magnetic mineral concentration and grain size (Chapter 14). It is now becoming apparent that geomagnetic intensity variations, as recorded in sediments, may well be the next important magnetostratigraphic tool, promising to provide high-resolution global correlation within polarity chrons. The authors thank the following people, who made important suggestions for improving the manuscript: B. Clement, D. Hodell, K. Huang, E. Irving, D. V. Kent, E. Lindsay, B. MacFadden, M. McElhinny, S. K. Runcorn, and J. S. Stoner. Parts of this book were written while one of us (J.E.T.C.) was on sabbatical in Kyoto (Japan) and Gif-sur-Yvette (France). M. Torii, C. Laj, and C. Kissel made these sabbaticals possible. Finally, we express special appreciation to Marjorie Opdyke, who patiently and skillfully produced the original figures for this book with loving attention to detail. The authors of Magnetic Stratigraphy received their doctoral training from the University of Newcastle-upon-Tyne, where Professor S. K. Runcorn inspired their life-long interest in paleomagnetism. We are saddened by his murder in December of 1995 and dedicate this volume to Dr. Runcorn as an expression of our gratitude to him as our friend and mentor.

Neil D. Opdyke James E. T. Channell

Table Key

Throughout this book, compilations of important magnetostratigraphic studies are presented as black/white (normal/reverse) interpretive polarity columns. These are accompanied by tables that include the following information for each study. Rock unit Age range Region h NSE NSI NSA M D

RM

DM AD

A designation of the rock unit, usually as a formation name, locality, or core/hole/site number Age range of the stratigraphic section(s) Country or region of study Site latitude (positive :north, negative :south) Site longitude (positive : east, negative :west) Number of sections Number of sites or site interval (m). P (includes passthrough magnetometer measurements) Number of samples per site Section thickness (m) Associated absolute age control: K (K-Ar), A (4~ 39Ar), F (fission track), R (Rb/Sr), S (strontium isotopic ratios), O (~180), As (astrochronology) Rock magnetic measurements/observations: K (susceptibility), J (Js/T), I (IRM), V (VRM), A (ARM), H (hysteresis measurements), TdI (thermal demagnetization of IRM), T (electron microscopy), O (optical microscopy), X (X.ray) Demagnetization procedure: A (alternating field), T (thermal), C (acid/chemical) How was the magnetization direction determined? B (blanket demagnetization), Z (Zijderveld/orthogonal ooo

XIII

xiv

Table Key

A

NMZ NCh %R RT

F.C.T. Q

projections), V (vector analysis), P ("principal component" or 3D least squares analysis) Data presentation: F (Fisher statistics presented), D (declination plotted), I (inclination plotted), V (VGP latitudes plotted) Number of magnetozones (polarity zones) Number of chrons Percentage of reverse polarity Reversal test: R+ [positive in the A, B, C rating of McFadden and McElhinny (1990) or according to McElhinny, (1964)], R - (negative) F (fold test), C (conglomerate test), Co (baked contact test); indicated as positive or negative Reliability index (see Chapter 5.7)

It should be noted that the quality of the data is sometimes not reflected in the reliability index. For example, some marine sediments are such good magnetic recorders that it is possible to get excellent quality magnetostratigraphic data with only basic techniques such as blanket demagnetization with alternating fields, as can be seen in some whole core data from the Ocean Drilling Program.

Introduction and History

1.1 Introduction Paleomagnetism has had profound effects on the development of Earth sciences in the last 25 years. In the early days, paleomagnetic studies of the different continental blocks contributed to the rejuvenation of the continental drift hypothesis and to the formation of the theory of plate tectonics. Paleomagnetism has led to a new type of stratigraphy based on the aperiodic reversal of polarity of the geomagnetic field, which is now known as magnetic polarity stratigraphy. Magnetic polarity stratigraphy is the ordering of sedimentary or igneous rock strata into intervals characterized by the direction of magnetization of the rocks, being either in the direction of the present Earth's field (normal polarity) or 180~ from the present field (reverse polarity). This new stratigraphy greatly influenced the subject of plate tectonics by providing the chronology for interpretation of oceanic magnetic anomalies (Vine and Matthews, 1963). In the last 25 years, the geomagnetic polarity time scale (GPTS) has become central to the calibration of geologic time. The bridge between biozonations and absolute ages and the interpolation between absolute ages are best accomplished, particularly for Cenozoic and Late Mesozoic time, through the GPTS. In most Late Jurassic-Quaternary time scales, magnetic anomaly profiles from ocean basins with more or less constant spreading rates provide the template for the GPTS. Magnetic polarity stratigraphy on land or in deep sea cores provides the link between the GPTS and biozonations/bioevents and hence geologic stage boundaries. Radiometric absolute ages are correlated either directly to the GPTS in magnetostratigraphic section or indirectly through

2

1 Introduction and History

biozonations, and absolute ages are then interpolated using the GPTS (oceanic magnetic anomaly) template. In view of the paramount importance of the calibration of geologic time to understanding the rates of geologic processes, the contribution of magnetic polarity stratigraphy to the Earth sciences becomes self-evident. The dipole nature of the main geomagnetic field means that polarity reversals are globally synchronous, with the process of reversal taking 103-104 yr. Magnetic polarity stratigraphy can therefore provide global stratigraphic time lines with this level of time resolution. Three other techniques in magnetic stratigraphy, not involving the record of geomagnetic polarity reversals, have become increasingly important. Rock magnetic stratigraphy refers to the use of nondirectional magnetic properties (such as magnetic susceptibility and laboratory-induced remanence intensities) as a means of stratigraphic correlation. Paleointensity magnetic stratigraphy, the use of the record of geomagnetic paleointensity, and secular variation magnetic stratigraphy, the use of secular directional changes of the geomagnetic field, have been used as a means of stratigraphic correlation in Quaternary sediments. Individual texts exist for paleomagnetism in general, such as those of Irving (1964), McElhinny (1973), Tarling (1983), Butler (1992), and Van der Voo (1993); however, none is available for the rapidly growing discipline of magnetic stratigraphy.

1.2 Early Developments Directions of natural remanent magnetization (NRM) reverse with respect to the present ambient magnetic field of the Earth were known to early pioneers in paleomagnetic research. Brunhes (1906) reported directions of magnetization in Pliocene lavas from France that yielded north-seeking magnetization directions directed to the south and up, rather than to the north and down. He attributed this behavior to a local anomaly of the geomagnetic field. Brunhes demonstrated that the baked contacts of igneous rocks are magnetized with the same polarity as the igneous rock. This was the first use of what is now referred to as the baked contact test, which can be used to determine the relative age of magnetizations in the vicinity of an igneous contact. When lava flows are extruded, or dikes are intruded into a host rock, the temperature of the rock or sediment into which the magma is intruded is raised above the blocking temperature of the magnetic minerals. The magnetization is, therefore, reset in the direction of the field prevailing at the time of baking. Brunhes carried out this test for both normally and reversely magnetized igneous contacts.

1.2 Early Developments

3

This work was followed by that of Matuyama (1929) on volcanic rocks from Japan and north China. This study included many examples of lavas that yielded directions that were reverse with respect to the present geomagnetic field (Fig. 1.1). Matuyama correctly attributed the reverse directions of magnetization to a reversal of the geomagnetic field. Matuyama separated his samples into two groups: Group I, which were all Pleistocene in age, gave NRM directions which grouped around the present direction of the geomagnetic field. The group II directions were antipodal to those of group I and to the present directions of the geomagnetic field, and most of these lavas were older than Pleistocene in age. This was the first hint that the polarity of the geomagnetic field might be age dependent. These early studies were followed by those of Chevallier (1925), Mercanton (1926), and Koenigsberger (1938). The modern era of studies of reversals of the geomagnetic field began with those of Hospers (1951, 1953-1954) in Iceland and Roche (1950, 1951, 1956) in the Massif Central of France, which elaborated on the early work of Brunhes. In both of these studies, the relative age and position of the rocks were fixed stratigraphically. As in Matuyama's studies, the results indicated that rocks and sediments designated as Upper Pleistocene and Quaternary in age possessed normal directions of magnetization, whereas

6

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SAISYUTO -Z )t MANCHURIA 1.1 Directions of natural remanent magnetization in basalts from China and Japan

(after Matuyama, 1929).

1 Introduction and History

reverse directions of magnetization appeared in rocks of early Pleistocene or Pliocene age. Einarsson and Sigurgeirsson (1955) utilized a magnetic compass to detect reversals in Icelandic lavas and began to map the distribution of magnetic polarity zones in the rock sequences. They found about equal thicknesses of normal and reverse polarity strata, which implied that the magnetic field had little polarity bias in the Late Cenozoic. A study by Opdyke and Runcorn (1956) on lava flows from the San Francisco Peaks of northern Arizona showed that all young lavas studied were normally magnetized, whereas the stratigraphically older series of flows contained reversely magnetized lavas. Based on these studies, it was assumed that the last reversal of the field took place close to the Plio-Pleistocene boundary. Rutten and Wensink (1960) and Wensink (1964) built on Hosper's studies in Iceland and demonstrated that the youngest lavas were normally magnetized and that older underlying lavas were reverse, and these in turn were underlain by a normal sequence. They subdivided the Plio-Pleistocene lavas into three magnetozones N1-R1-N2. They correlated N2 to the Astien and R1 to the Villafranchian and deduced that the earliest glaciations in Iceland were of Pliocene age. Magnetometers capable of measuring the magnetization of igneous rocks had been available since before World War II. Johnson et al. (1948) developed a rock generator magnetometer which could measure natural remanence in sedimentary rocks. Blackett (1956) developed astatic magnetometers which were a further improvement in sensitivity. Paleomagnetic study of sedimentary rocks began with the classic study of Creer et al. (1954) which documented 16 zones of alternating polarity in a 3000-m section of the Torridonian sandstone (Scotland). These rocks passed the fold test, indicating a prefolding magnetization, implying that reversals were a long-term feature of the geomagnetic field. These authors also reported reversals from rocks of Devonian and Triassic age. Simultaneous with the developments outlined above, which seemed to favor reversal of the dipole geomagnetic field as an explanation for the observations, a discovery was made in Japan which cast doubt on the field reversal hypothesis. Nagata (1952), Nagata et al. (1957), and Uyeda (1958) demonstrated that samples from the Haruna dacite, a hyperthermic dacite pumice from the flanks of an extinct volcano in the Kwa district of Japan, was self-reversing. The striking magnetic property of this rock is that, when heated to a temperature above 210~ and cooled in a weak magnetic field, the acquired thermal remanent magnetization (TRM) is in opposition to the applied field. Since self-reversal in rocks could be demonstrated, it became important to ascertain whether all reverse directions could be explained in this way, or whether both the self-reversal phenomenon and reversal of the main dipole field were taking place.

1.3 Evidence for Field Reversal

1.3 Evidence for Field Reversal The baked contact test first employed by Brunhes played an important role is settling the self-reversal/field reversal controversy. If field reversal is the norm, one would expect baked contacts to be magnetized in the same polarity as the baking igneous rock. If self-reversal is the norm, one would expect the baked contacts to have different polarity from the igneous rock in a significant proportion of the cases, since it is unlikely that both igneous and country rock would have the same magnetic mineralogy. Wilson (1962) conducted a survey of such studies and found that in 97% of the contacts studied, the polarity of the baked contact was the same as that of the igneous body. These observations supported the hypothesis of geomagnetic field reversal. By the late 1950s, several lines of evidence supported the hypothesis that the geomagnetic field reverse polarity. (1) All rocks of Late Pleistocene and recent age, both igneous and sedimentary, that had been studied possessed normal polarity magnetizations. (2) Reverse polarity magnetizations were observed in rocks of Lower Pleistocene or older age and in sedimentary and igneous rocks involving different magnetization processes. (3) Transitional (intermediate) directions were observed in Iceland in both reverse-to-normal and normal-to-reverse transitions, as well as in the Torridonian of Scotland. (4) Self-reversal had been observed in the laboratory in only one rock type (Haruna dacite) although many attempts had been made to replicate the experiment in other rock types. (5) As described above, in 97% of the contact tests, the baked contact and the baking rock possessed the same polarity. By the late 1950s, many paleomagnetists believed in geomagnetic field reversal although geophysicists in general were still skeptical. In the early 1960s, the stage was set for the convincing demonstration of geomagnetic field reversal. A series of interdisciplinary studies were carried out in which K/Ar dating and the measurement of magnetization polarity were carried out on the same lavas (Cox et al., 1963; McDougall and Tarling, 1963b). These studies established that rocks of the same age had the same polarity and led to the establishment of the first radiometrically dated polarity time scale. As information accumulated throughout the 1960s, the magnetic polarity sequence was progressively modified (Fig. 1.2), particularly with the discovery of short magnetic polarity intervals (Gromm6 and Hay, 1963). The long intervals of constant polarity of the geomagnetic field were designated by Doell and Dalrymple (1966) as magnetic epochs and named after the pioneers in the study of the geomagnetism (Brunhes, Matuyama, Gauss, and Gilbert); the shorter intervals (events)

1 Introduction and History

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Time (Ma) Figure 1.2 Evolution of the geomagnetic polarity time scale of the last 2.6 My (after Watkins, 1972).

were named after the locality of discovery such as Jaramillo Creek in New Mexico and Olduvai Gorge in Tanzania. It should be noted that the time scales were produced from lavas that were scattered over several continents

1.3 Evidence for Field Reversal

7

and oceanic islands and were not in stratigraphic superposition, representing a major departure from standard stratigraphic procedures. The anomalous situation arose in which, unlike classical stratigraphy, no type sections for the magnetic epochs (polarity chrons) are available, yet type localities for magnetic events (polarity subchrons) are known. The developments which led to the first dated polarity time scale were soon followed by studies of magnetic stratigraphy in sediments and investigations of oceanic magnetic anomalies. Khramov (1958) in the Soviet Union reported reverse and normal magnetizations in Plio-Pleistocene nonmarine sedimentary sequences from the Caucasus. These studies were followed by the discovery of reversals in marine sediments (Harrison and Funnel, 1964; Linkova, 1965). Opdyke et aL (1966) observed the same sequence of polarity

V e m a core 21 - 148 Inclination (~ -60

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Eltanin track 19

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Figure 1.3 Comparison of oceanic magnetic anomalies from ship track Eltanin 19 (Pitman

and Heirtzler, 1966) with magnetic polarity stratigraphy in core V21-148 (Opdyke, 1968) and the magnetic polarity time scale of Doell and Dalrymple (1966).

8

1 Introduction and History

reversals in marine sediments that had been compiled by Cox et al. (1963) from dispersed volcanic outcrops. The same reversal sequence was then observed in oceanic magnetic anomalies by Vine and Wilson (1965) and Pitman and Heirtzler (1966). The fact that the same reversal sequence appeared to have been recorded by three different recording media (Fig. 1.3) closed the argument in favor of field reversal and relegated self-reversal to a subsidiary process. The resolution of this argument in favor of field reversal tipped the scales in favor of the hotly debated hypothesis of plate tectonics and sea floor spreading.

The Earth's Magnetic Field

2.1 Introduction The magnetic field of the Earth has fascinated human beings for well over 2 millennia. The Chinese invented the magnetic compass in the second century B.C. (Needham, 1962) and knowledge of the magnetic compass reached Western Europe over a thousand years later in the twelfth century A.D. The first truly scientific paper on geomagnetism was written in 1262 by Petrus Peregrinus and entitled "Epistola de Magnete" (Smith, 1970). In a series of experiments on lodestone (magnetite) spheres, Peregrinus defined the concept of polarity and defined the dipolar nature of a magnet stating that like poles repel and unlike poles attract; however, the treatise was not published until 1558. The work by Peregrinus undoubtedly influenced the later work of William Gilbert, who published the book "De Magnete" in 1600. Gilbert studied the variation of the angle of inclination (or magnetic dip) over lodestone cut into the shape of a sphere, perhaps one of the greatest model experiments ever made in Earth science. His conclusion was that "the Earth itself is a great magnet." From the 14th century to the present, compasses were widely used as a navigation aid on both land and sea and were carried by explorers throughout the world. The angular distance between the polar star, which does not move in the heavens, and the direction of the north-seeking end of the compass needle was often faithfully recorded. This magnetic deviation from true north is called the magnetic declination (D). In midlatitudes, from 40~ to 40~ the declination of the present geomagnetic field can vary up to 40 ~ from true north. In polar regions, near the magnetic poles, declination anomalies of up to 180 ~ are observed (Fig. 2.1).

10

2 The Earth's Magnetic Field 80~

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Figure 2.1 Declinationfor the International Geomagnetic Reference Field (IGRF) 1990(after Baldwin and Langel, 1993).

If a magnetized needle is suspended on a fiber and allowed to swing freely, the north-seeking end of the needle will not only point north, it will also point down in the northern hemisphere and up in the southern hemisphere. This property of the earth's field, its inclination, was discovered by George Hartmann in 1544 (Smith, 1968). If a dip needle which measures the angle of magnetic inclination is carried from high northern latitudes to high southern latitudes, it will be seen that the inclination of the field varies systematically from vertical down at the north magnetic pole, to the horizontal at some point at low latitude (magnetic equator), and vertical once more (but with the north-seeking end of the needle pointing vertically upward) at the south magnetic pole (Fig. 2.2). It should be noted that the north and south magnetic poles are not 180~ apart and that the magnetic equator is not equidistant from the two magnetic poles. If the geomagnetic field is analogous to that of the magnetized spheres studied by William Gilbert, the intensity of the field, as well as the inclination and declination, would vary systematically over the Earth. After Gauss invented the deflection magnetometer, this was found to be the case, the magnetic field strength at the magnetic poles being almost twice that at the magnetic equator. The field at any point on the Earth's surface is a vector (F) which possesses a component in the horizontal plane called the horizontal component (H) which makes an angle (D) with the geographical meridian (Fig.

2.1 Introduction

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Figure 2.2 Inclination for the International Geomagnetic Reference Field (IGRF) 1990 (after Baldwin and Langel, 1993).

2.3). The declination (D) is an angle from north measured eastward ranging from 0 ~ to 360 ~ The inclination (I) is the angle made by the magnetic vector with the horizontal. By convention, it is positive if the north-seeking vector points below the horizontal or negative if it points above. The horizontal component (H) (Fig. 2.3) has two components, one to the north X and one to the east Y. The following equations relate the various quantities: H = F cos I

Z = Fsin I

Tan I = Z / H

(2.1)

X=

Y= HsinD

Tan D = Y / X

(2.2)

HcosD

F 2 = H 2 + 1 2 __ X 2 _+_ y 2 + Z 2.

(2.3)

The dipole nature of the geomagnetic field, originally suggested by Gilbert, was first put to a mathematical analysis by Gauss (1838), who used measurements of declination, inclination, and intensity from 84 widely spaced locations to derive the first 24 coefficients in the spherical harmonic expansion of the geomagnetic field. The results of this analysis indicated that, within the uncertainties of the data, the geomagnetic field had no sources external to the Earth and that it had a dominant dipole component. Since 1835, when Gauss first determined the intensity of the dipole moment, the calculations have been repeated many times, and this value has been decreasing steadily over the last century and a half. During this century,

12

2 The Earth's Magnetic Field

//North (Geographical)Magneti c meridian X .....

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Downwards Figure 2.3 The direction

and intensity of the total field vector (F) resolved into declination from geographical north (D) and inclination from horizontal (I). Equations (2.1), (2.2), and (2.3) relate the various quantities.

the decay rate has increased, and during the last 30 years has reached a rate of 5.8%/century (see Barton, 1989). If the trend were to continue, the main dipole field would disappear by the year 4000 A.D., a possible but unlikely event. Modern spherical harmonic analyses confirm Gauss' findings and lead to the conclusion that: (1) the geomagnetic field is almost entirely of internal origin, (2) ---90% of the field observed on the surface can be explained by a dipole inclined to the Earth's axis of rotation by 11.5~ (Fig. 2.4), (3) the magnitude of the dipole moment is 7.8 • 10 22 A m 2, (4) the axis of the centered dipole (geomagnetic pole) emerges in the northern hemisphere between Greenland and Ellesmere Island at 79.0~ 70.9~ The great circle midway between the geomagnetic poles is called the geomagnetic equator (Fig. 2.4). Magnetic observations of the elements of the geomagnetic field were begun as early as the 16th century in London and the 17th century in Paris. It was realized early on that the declination and inclination of the

13

2.1 Introduction

Iraphic pole

Geomagnetic north North magnetic pole (1=90~ j

1.5 ~

Magnetic equator (1=0 Geographic equatoT" """

c_

Geomagnetic equator

"~--Best fitting dipole South magnetic pole (1=-90)/

Geographic pole

Geomagnetic south pole

Figure 2.4 Graphical presentation of the magnetic, geomagnetic, and geographic poles and equators (after McElhinny, 1973).

geomagnetic field change with time (Fig. 2.5). This relatively rapid change of the magnetic field with time is called secular variation. The declination of the field at London has been changing at a rate of about 14~ per century from 0 ~ in 1650 to 24~ in 1800 and is now 5~ The change in the magnetic elements is not constant over the Earth's surface and, for example, changes less rapidly over the Pacific Ocean than in, say, South America. Global secular variation is often displayed as "isoporic" charts, which are contour maps of equal annual change of a particular element of the field, such as the vertical component. From isoporic charts constructed from observations dating back about 200 years, isoporic centers have drifted westward at rates of up to about 0.28~ over this time interval. This observation is a manifestation of the westward drift of dipole, quadrupole, and octupole components of the field. Drift rates of these components can be calculated directly from the spherical harmonic coefficients and their changes with time. Changes in the rate of westward drift and the growth and decay of isoporic features occur on a decadal time scale. The westward drift is generally attributed to differential angular velocity of Earth's lithosphere and outer core, where the nondipole components of the field are thought

14

2 The Earth's Magnetic Field

64

(/) 13) (!)

~ .

co

' I I';~---~. D'Paris 187y 850J,L 1913 ,London

66 -

68

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Declinations (degrees) Figure :2.5 Change of declination and inclination at London and Paris from observatory records (after Gaiber-Puertas, 1953).

to originate. It is important to note that although westward drift is dominant, not all secular variation features drift westward; some are more or less stationary and some are moving slowly eastward. The geomagnetic field varies over a wide range of time scales (Table 2.1). The longer term behaviors (4-7, Table 2.1) are produced in the Earth's interior, and the short-term behaviors (1-3, Table 2.1) are atmospheric or ionospheric in origin. Secular variations of the field, first observed in observatory records, can be monitored further back in time using paleomagnetic data from marine and lake sediments with high accumulation rates. The longer term variations (5-7, Table 2.1) have also been documented by paleomagnetic study of rocks and sediments. Since some lavas and archaeological artifacts can become permanently magnetized in a matter

Table 2.1 Scales of Geomagnetic Variability Geomagnetic behavior 1. 2. 3. 4. 5. 6. 7.

Pulsations or short-term fluctuation Daily magnetic variations Magnetic storms Secular variations Magnetic excursions Reversal transition Interval between reversals

Duration minutes hours hours to days 102-103 yr 103-104 yr 103-104 yr 105-106 yr

2.2 The Dipole Hypothesis

|

of hours, it is theoretically possible to detect ancient field behavior on this time scale from paleomagnetic data.

2.2 The Dipole Hypothesis The fact that the geomagnetic field can be modeled as a geocentric dipole is of fundamental importance in paleomagnetic studies. If we assume that the time-averaged dipole is aligned with the axis of rotation of the Earth and situated at the center of the Earth (an axial geocentric dipole), the geomagnetic and geographic axes will coincide, and the geomagnetic equator will coincide with the geographic equator. The geomagnetic inclination (I) will vary systematically with latitude (A) according to the formula: Tan (I) = 2 Tan (A). If M is the moment of the geocentric dipole and the radius of the Earth is r, the horizontal (H) and vertical (Z) components at any latitude (A) can be derived from the following geometric relationship (Fig. 2.6), where I is the inclination from the horizontal. Using the formulae used by Gauss for his deflection magnetometer, H= (McosA)/r ~ Z = (2M sin A ) / r ~

(2.4)

N

Figure 2.6

Field lines at the Earth's surface for the axial geocentric dipole (after McE1hinny, 1973).

|6

2 The Earth's Magnetic Field

Tan I = Z/H = 2 Tan (A),

(2.5)

if D = 0 ~ colatitude (0) is given by TanI=

2cos 0 ( 0 <

0 < 180~

(2.6)

The relationship makes it possible to determine the paleolatitude of a site from mean paleomagnetic directions at any point on the Earth's surface. In the discussion given above, the declination (D) was presumed to be zero, although in most paleomagnetic investigations this is not so. The paleomagnetic pole represents the position on the Earth's surface where the dipole axis (oriented to give a field in agreement with the mean paleomagnetic direction) cuts the surface of the globe. Every paleomagnetic investigation, if successful, will yield a declination (D) which represents the angle between the paleomagnetic vector and the present geographic meridian. The declination defines a great circle that passes through the sampling site and the paleomagnetic pole. The mean paleomagnetic inclination gives the distance along the great circle (or paleomeridian) to the paleomagnetic pole, according to the dipole formula given above (2.5). The pole position can be calculated in the present geographical coordinate system using the following relationship. If Dm, Im are known at a sampling site with latitude A and longitude 9 then the position of the paleomagnetic pole P (A',O') can be calculated sin A' = sin A cos 0 + cos A sin 0 cos Dm

(2.7)

~' =~+/3

(2.8)

(when c o s 0 > s i n A s i n A ' )

~' = ~ + (180-/3)

(when c o s 0 < s i n A s i n A ' ) ,

(2.9)

where sin/3 = sin 0 sin D/cos A'. By convention, latitudes are positive for the northern hemisphere and negative for the southern hemisphere; longitudes are measured eastward from the Greenwich meridian and lie between 0 and 360 ~ Other symbols are defined in Fig. 2.7. In paleomagnetic research, two types of paleomagnetic pole positions are calculated. A virtual geomagnetic pole (VGP) is a pole calculated from magnetic data which represent a very short period of time, such as the observatory record of D and I of the present field, or a spot paleomagnetic field reading from a single lava flow. A paleomagnetic pole, on the other hand, represents the field averaged over a considerable length of time (on the order of 50 ky) such that short-term (secular) variations are averaged out. A paleomagnetic pole position is essentially the mean vector sum of many virtual geomagnetic pole positions. In the middle 1950s, it was known that the mean Plio-Pleistocene paleomagnetic poles more or less coincide with the pole of rotation of the Earth. A compendium of poles by Tarling (1983) yields an average

2.3 Models of Field Reversal

]7

N

Figure 2.7 Calculation of a paleomagnetic pole from mean directions of magnetization (Dm, Im) at sampling site (S) generates a paleomagnetic pole (P) (after McElhinny, 1973).

paleomagnetic pole indistinguishable from the present pole of rotation (Fig. 2.8). These observations support one of the basic hypotheses in paleomagnetism, the dipole hypothesis, which states that the Earth's dipole field averages to an axial geocentric dipole over periods of time longer than the averaging time for secular variation (10-50 ky). The dipole hypothesis may be tested using widely spaced sampling sites for rocks of the same age, or by testing the relationship of inclination to site latitude in marine cores. Opdyke and Henry (1969) and Schneider and Kent (1990a) demonstrated that Pleistocene paleomagnetic inclinations in widely spaced marine cores are consistent with the dipole hypothesis (Fig. 2.9). Merrill et al. (1979) and Schneider and Kent (1988, 1990a) have documented asymmetry between normal and reverse intervals over the last few million years, which can be accounted for by increasing the quadrupole term for reverse polarity intervals relative to normal polarity intervals. This asymmetry is difficult to explain due to the symmetry of the induction equations (Gubbins, 1994). Outer core boundary conditions (in the mantle or inner core) can be invoked to break the symmetry of the dynamo equations (see Clement and Stixrude, 1995).

2.3 Models of Field Reversal Although the geomagnetic field may be molded as if the Earth were a permanent magnet, it is clear that this is not the case as thermal disordering (Curie) temperatures of permanently magnetized materials are exceeded

|8

2 The Earth's Magnetic Field

9

o

Figure 2.8 Paleomagnetic poles for igneous rocks younger than 20 Ma (after Tarling, 1983).

60 ~ r C)

30 ~

,,m

C~

O

tim t~ t,-

0

m

m

.C

S

-60 ~ -30 ~

0~

30 ~

60 ~

N

Latitude Figure 2.9 Inclination versus latitude in Pleistocene marine sediments. Line is expected inclination derived from the dipole formula (after Schneider and Kent, 1990a).

2.3 Models of Field Reversal

|

by burial greater than 10-20 km. The recognition of polarity reversals of the dipole field also provided a powerful argument against permanent magnetization. In 1947, Blackett proposed that planetary dipole fields are an intrinsic property of rotating bodies and subsequently designed experiments that refuted the claim (Blackett, 1952). Beginning with the work of Elsasser (1946) and Bullard (1949), it is now known that the geomagnetic field arises from convective and Coriolis motions in the fluid nickel-iron outer core. According to dynamo theory, outer core flow can amplify a small instantaneous axial field to observed dipole values, although the details of the process remain poorly constrained. No detailed theory predicting the form of geomagnetic secular variation or the mechanism of geomagnetic reversals is yet available. Gubbins (1994) and Jacobs (1994) have summarized recent developments on this topic. Uncertainties in the details of the geomagnetic field generation process lead to little consensus on the process of geomagnetic polarity reversal. The models of Cox (1968, 1969), Parker (1969), and Levy (1972) involved interaction between the main dipole field and the nondipole field. According to Parker (1969) and Levy (1972), changes in distribution of cyclonic eddies, which account for the nondipole field, can produce regions of reverse torodial field which effect reversal of the poloidal (dipole) field. According to Cox (1968, 1969), the critical factor is the relative field strength of the high-frequency oscillations of the nondipole component relative to the lower frequency (cyclic) oscillations of the dipole component. A highamplitude nondipole field coincident with a low in the main dipole field could, in this model, produce an instantaneous reverse total field, which is then reinforced by the dynamo process. According to Laj et al. (1979), the supposed coupling of the nondipole and dipole fields in the Cox model is unlikely in view of the different time constants of the two field components, and the mechanism (if feasible) would be expected to generate very frequent reversals. Hillhouse and Cox (1976) supposed that the dipole component of the field dies in one direction and grows in the opposite direction during a polarity reversal and that the nondipole components do not change over times sufficient to span several reversals. The model predicts that sequential reversals, of opposite and same sense, recorded at a sampling site should yield similar transitional VGP paths. This has not been found to be the case and hence this model is no longer favored. The models of Hoffman (1977) and Fuller et al. (1979) are based on the idea that the reversal process floods through the core from a localized zone. These models can be tested as they predict the VGP paths during a polarity transition for either quadrupole or octupole transition field geometries. The VGP paths are predicted to lie either along the site longitude

20

2 The Earth's Magnetic Field

(near-sided) or 180 ~ away (far-sided) depending on the sense of the reversal, the hemisphere of observation, and where in the core the reversal process is initiated. These models are generally no longer favored due to the lack of evidence for near-sided or far-sided VGP paths. Olson (1983) attributed reversal to a change in sign of the helicity (the correlation between turbulent velocity and vorticity), triggered by a change in the balance between heat loss at the core-mantle boundary and solidification at the inner-outer core boundary. In this model, the duration of a polarity chron will depend on the strength of the dipole field and the process of reversal would be rapid (---7500 yr), more or less consistent with paleomagnetic observations. The solar dipole fluid reverses polarity approximately every 11 years (Eddy, 1976) and the observations of the solar reversal process provide clues to the reversal process on Earth (see Gubbins, 1994). The high frequency of solar field reversal allows the process r o b e observed directly. In addition, the flux distribution at the solar surface can be monitored, in contrast to the Earth, where the flux distribution at the core-mantle boundary must be approximated by downward continuation. Relatively low temperature sunspots are regions of intense magnetic field, considered as sites of expulsion of toroidal flux. The locations of sunspot pairs move from the solar poles toward the equator, leading up to a reversal of the solar dipole field. Bloxham and Gubbins (1985) and Gubbins (1987a,b) interpret flux concentrations in maps of the radial component of the geomagnetic field as due to eddies at the core-mantle boundary, which they liken to sunspots. Two "core spots" have been identified beneath South America and southern Africa. The process at the Earth's core-mantle boundary leading to the reversal of the geomagnetic dipole may be similar to the solar reversal process, with flux spots beneath South America and southern Africa being analogous to sunspots. These core spots produce local fields opposed to the Earth's main dipole field. Bloxham and Gubbins (1985) have suggested that these areas of high magnetic flux account for the high rate of secular variation in the Atlantic region. The core spots are apparently strengthening as the main dipole field decays. They are apparently moving to the south and could possibly lead to a reversal of the main dipole field. As the core spots in these regions become larger, they might be observed at the surface as large regional deviations of the field (magnetic excursions).

2.4 Polarity Transition Records and VGP Paths In the past decade, an increasing amount of high-quality paleomagnetic data have been acquired from polarity transitions in high deposition rate

2.4 Polarity Transition Records and VGP Paths

~|

sedimentary sequences (e.g., Clement, 1991; Tric et al., 1991a; Laj et al., 1991) and in volcanic rocks (e.g., Prdvot et al., 1985; Hoffman, 1991, 1992). The records provide important insights into the configuration of the field during a polarity transition; however, the interpretations are controversial. Some data indicate that transition fields have a large dipolar component. For example, the records of the Cobb Mountain subchron from the western Pacific and North Atlantic have similar VGP (virtual geomagnetic pole) paths, implying dipolar transitional fields (Clement, 1992). The VGP paths of some spatially distributed Matuyama-Brunhes polarity transitions coincide (Clement, 1991) whereas others do not (Valet et al., 1989). Another controversy has involved the apparent clustering of polarity transition VGPs along two longitudinally constrained paths, one at -80~ longitude over North and South America and the other at -100~ over eastern Asia and Australia (Fig. 2.10). Some studies emphasize the statistical significance of the preferred VGP paths (e.g., Clement, 1991; Laj et al., 1991, 1992). Valet et al. (1992) have disputed both the claim that transition fields are dipolar and the statistical clustering of transitional VGPs to longitudinal paths. Quidelleur and Valet (1994) considered that the apparent clustering of VGP paths is biased by uneven site distribution. The highest concentration of sampling sites is in southern Europe, - 9 0 ~ away from the VGP longitudinal path across the Americas. Egbert (1992) showed that, except in the case where the transition field is entirely dipolar, the VGPs are likely to cluster --~90~ in longitude away from the sampling site. According to Constable (1992), the preferred VGP path across the Americas is apparent even after the sites located ---90~ away from the path are removed from the record. It is important to realize that preferred longitudinal VGP paths do not necessarily mean that the geomagnetic field is dipolar during polarity transitions (Gubbins and Coe, 1993). Prdvot and Camps (1993) examined a large population of transitional records from volcanic rocks younger than middle Miocene and concluded that the transitional VGPs have no statistical bias in their distribution. On the other hand, more selective use of data from volcanic rocks leads to a different conclusion. Hoffman (1991, 1992) found that transitional VGPs from distributed sites of volcanic rocks younger than 10 Ma fall into two clusters in the southern hemisphere, one over Australia and the other over South America (Fig. 2.10). The clusters are located close to the Asian and American longitudinal VGP paths resolved from sedimentary transition records. These data lead to the conclusion that transitional field states are long-lived (i.e., characteristic of sequential reversals) and have a strong dipolar component, which is inclined to the Earth's rotation axis. Bloxham and Gubbins (1987) have downward continued the field since 1700 A.D. to the core-mantle boundary, and it appears that magnetic flux

22

2 The Earth's Magnetic Field

Figure 2.10 (a) Clustered VGPs concentrated over Southeast Asia and South America/Antarctica from polarity transition records from lava flows. (b) Means of VGPs from lava flows superimposed on shaded preferred longitudinal VGP paths from sedimentary reversal transition records (after Hoffman, 1992).

at the core-mantle boundary is also concentrated approximately along the east Asian and American longitudinal paths. Gubbins and Kelly (1993) have analyzed paleomagnetic data for the last 2.5 My from both lavas and sediments and suggested that the magnetic field at the core's surface has remained similar to that of today over this period of time. The persistence of the preferred paths through time and the fact that paths pass through regions of high modern flux activity at the core-mantle boundary suggest that these regions are pinned in their present position and have been for some time. The long time scales involved indicate that the position of these regions is controlled by mantle (as opposed to outer core) processes. Laj et al. (1991) and Hoffman (1992) have pointed out that radial flux centers

2.4 Polarity Transition Records and VGP Paths

~3

associated with rapid secular change in today's field are closely associated with regions of fast-P-wave propagation in the lower mantle (Dziewonski and Woodhouse, 1987), indicating the presence of relatively cold mantle; however, this interpretation of mantle tomography in terms of temperature variations is not universally accepted. The preferred longitudinal VGP paths lie either side of the Pacific Ocean, which is a region characterized by low-amplitude secular variation. At present, there are no loci of rapid secular change in the Pacific region and none have been known since measurements of the magnetic field in the region began several hundred years ago. Runcorn (1992) suggested that this low rate of secular change may be caused by a well-developed high conducting D" shell under the Pacific. According to Runcorn (1992), as the dipole begins to reverse polarity, the magnetic field induces currents in the Pacific D" shell, causing a torque that rotates the outer core relative to the mantle until the reversing dipole lies along the American or East Asian longitudinal paths. Clement and Stixrude (1995) have invoked magnetic susceptibility anisotropy of the inner core to account for several problematic features of the paleomagnetic field. The Earth's inner core exhibits anisotropy in seismic velocity which can be explained in terms of the preferred orientation of hexagonal close-packed (hcp) iron. Room temperature analogs of hcp iron have a strong magnetic susceptibility anisotropy, which could produce an inner core field inclined to the rotation axis. This may help to explain not only the well-known far-sided effect (Wilson, 1972) but also the preferred VGP paths during polarity transitions. When the outer core field is weak during a polarity reversal, the persisting inner core field may bias the total field to produce the observed VGP paths and clusters observed in transition records. During polarity reversal, as the outer core field grows in the opposite direction, it may take a few thousand years to diffuse through the inner core and reset the inner core field. Polarity reversals in the paleomagnetic record appear to be accompanied by geomagnetic intensity lows (e.g., Opdyke et al., 1966), although in some sedimentary records, the apparent lows may be an artifact of the remanence acquisition process. Paleointensity studies by Valet and Meynadier (1993) imply reducing geomagnetic field intensity during PlioPleistocene polarity chrons, with intensity recovery immediately post reversal. The present axial dipole moment is 7.8 x ] 0 22 A m 2. If it is assumed that the calibration of the relative paleointensity to absolute values is correct, then the dipole moment during reversals falls to about 1 • 10 22 A m 2, or about a 10% of its present value. Valet and Meynadier (1993) pointed out that the lows in apparent paleointensity within the Brunhes Chron correspond to the age of documented field excursions (Ch. 14).

24

2 The Earth's Magnetic Field

2.5 Statistical Structure of the Geomagnetic Polarity Pattern As the GPTS is refined and our record of polarity reversals becomes more complete, estimates of the statistical structure of the polarity sequence become more realistic. From present to 80 Ma, since the Cretaceous long normal superchron, the polarity sequence approximates to a gamma process where the gamma function (k) equals ---1.72 (Phillips, 1977) or, in an updated estimate, ---1.5 (McFadden, 1989). The Poisson process is the special case of a gamma process where k = 1, corresponding to a random process. The observation that the gamma function is greater than 1 means that either the process has "memory," which reduces the likelihood of a reversal immediately post reversal, or that the record of reversals for the last 80 My is incomplete with many short-duration polarity chrons missing from the record. The oceanic magnetic anomaly record is characterized by numerous "tiny wiggles" referred to as cryptochrons in the time scale (Cande and Kent, 1992b). Inclusion of cryptochrons in the record can, depending on their duration, reduce the value of k to values very close to 1, indicative of a Poisson distribution (see Lowrie and Kent, 1983). Although the issue of tiny wiggles in the oceanic magnetic anomaly records is not settled, it is now popular to interpret them as due to paleointensity variations rather than short polarity chrons. If this is correct, it implies that polarity reversal of the geomagnetic field is not a Poisson process but that the probability of a reversal drops to zero immediately after a reversal and then rises to a steady value. The rate of reversal has varied with time and the variation in rate is quite smooth, with a steady decrease from 165 Ma to 120 Ma and a steady increase from 80 Ma to the present. The change in reversal rate results in a change in the potential resolution of magnetic polarity stratigraphy as a correlation tool. From 120 Ma to 80 Ma, during the Cretaceous normal superchron, the reversal process was not functioning, which may imply that the geomagnetic field is more stable in the normal polarity state. McFadden and Merrill (1984) analyzed the relative stability of normal and reverse polarity states and concluded that there is no evidence for a stability bias between the two polarity states. The existence of a superchron of reverse polarity in the Late Carboniferous-Permian of comparable duration to the Cretaceous normal superchron tends to support the notion that the polarity of superchrons is not dictated by stability bias. There is evidence, however, for a difference in the relative proportions of zonal harmonics which make up the time-averaged field for the two polarity states. On the basis of data for the last 5 My, Merrill and McElhinny (1977) concluded that the reverse polarity field has a larger quadrupole and octupole content than the normal

2.5 Statistical Structure of the Geomagnetic Polarity Pattern

25

polarity field. Schneider and Kent (1988) reached a similar conclusion based on the analysis of deep-sea core inclination data. This apparent asymmetry for normal and reverse polarity states may be attributed to an axis of magnetic susceptibility anisotropy in the inner core inclined to the rotation axis (Clement and Stixrude, 1995).

Magnetization Processes and Magnetic Properties of Sediments

3.1 Basic Principle All matter responds to an applied magnetic field, due to the effect of the field on electron motions in atoms, although the response is very weak for most materials. The ions of some transitional metals, notably Fe 2+, Fe 3+, and Mn 2+, carry an intrinsic spin magnetic moment and are referred to as paramagnetic ions. Materials which contain these cations exhibit an enhanced response to an external field. The weakest response is called diamagnetism. Diamagnetic materials do not contain paramagnetic ions in appreciable concentrations and when placed in a magnetic field (H), a weak magnetization (Mi) is induced, by the effect of the magnetic field on orbiting electrons, which is antiparallel to the applied field (H). The susceptibility (k) (where Mi - kH) is therefore small and negative. No (remanent) magnetization (Mr) remains after the magnetizing field is removed. Many of the most common minerals (such as quartz, halite, kaolinite, and calcite) are diamagnetic, and rocks which are composed largely of these minerals will have negative susceptibility. An enhanced response to a magnetic field is referred to as paramagnetism. Paramagnetic materials contain paramagnetic ions. The susceptibility is large (relative to diamagnetic susceptibility) and positive and due to the alignment of the intrinsic magnetic moments in the applied field. The susceptibility is proportional to the reciprocal of absolute temperature. When the.applied field is removed, the weak alignment of spin moments is randomized by thermal vibrations and paramagnetic materials do not 26

3.1 Basic Principle

27

retain a remanent magnetization. Many clay minerals and common ironbearing minerals such as siderite, ilmenite, biotite, and pyrite are paramagnetic. A few minerals, some iron oxides, oxyhydroxides, and sulfides, exhibit a third type of magnetic behavior called ferromagnetism (sensu lato). In these materials, the paramagnetic cations are closely juxtaposed in the crystal lattice such that spin moments are coupled either by direct exchange or by superexchange through an intermediate anion. Ferromagnetic susceptibility values vary widely but are high (relative to paramagnetic susceptibility values) and positive. Unlike diamagnetic and paramagnetic susceptibility, ferromagnetic susceptibility is strongly dependent on the strength of the applied field (H). At low H, the induced magnetization (Mi) is proportional to H, and therefore the initial susceptibility (k) is constant, and the magnetization is lost after removal of the applied field. However, at higher values of H, part of the magnetization is retained by the material after the removal of the applied field. This behavior can be summarized by use of a hysteresis loop, which is a plot of magnetization (M) against applied field (H) (Fig. 3.1). At low H, M increases proportionally (hence, k is constant for low H) and the magnetization process is reversible. At higher H, the magnetization is not proportional to H, the curve is not reversible, and a small proportion of the magnetization, referred to as the remanent magnetization (Mr), is retained by the material after field removal. Eventually, as

8 0 -

,

,

,

I

,

,

,

I

,

,

,

I

.

.

,

c

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I. ,

,

,

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,

,

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6O E

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4O 20 0 -20

Hc

C~

-40 -60 - 8 0

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-800

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.

-600

.

.

.

.

-400

.

.

.

.

-200

.

.

.

0

I

200

'

'

'

I

400

'

'

'

I

600

'

'

'

800

H: Applied Field (mT)

Figure 3.1 Hysteresisloop from a pink pelagiclimestone,illustratingthe saturation magnetization (Ms), saturation remanence (Mrs), and coerciveforce (Hc). The loop is constricted (waspwaisted) due to presence of low-coercivity magnetite and high-coercivity hematite.

28

3 Magnetization Processes and Magnetic Properties of Sediments

H increases further, the saturation magnetization (Ms) is reached. Removal of the applied field leaves the sample with a saturation remanent magnetization (Mrs). If the applied field is now increased in the opposite sense ( - H ) , the total magnetization of the sample will fall to zero at a value of - H referred to as the coercive force (Hc) for that sample. The back field which leaves the sample with zero remanent magnetization is referred to as the remanent coercivity or the coercivity of remanence (Hcr). The hysteresis loop can be completed by taking the sample to its saturation magnetization in the back field and imposing an increasing positive H until the saturation is achieved again. The hysteresis properties of ferromagnetic materials are destroyed at a characteristic temperature for the mineral referred to as the Curie or NEel temperature, above which the behavior is paramagnetic. The term ferromagnetism (senso lato) is used to refer to materials which show hysteresis behavior. There are various types of ferromagnetism which differ in magnetic properties due to different arrangements of spin moments within the crystal lattice. In the case of ferromagnetism (sensu stricto), the adjacent spin moments are parallel and of uniform magnitude (Fig. 3.2). This behavior is restricted to native iron, nickel, and cobalt where paramagnetic ions are sufficiently closely juxtaposed for direct exchange interactions. Minerals (on Earth) do not exhibit this type of behavior. In ferrimagnetism adjacent spin moments are antiparallel but a net moment results due to their unequal magnitude. In antiferromagnetism, adjacent spins are antiparallel and of uniform magnitude and hence no net moment results. If, however, the antiparallelism of the spin moments is imperfect due to canting, a weak net moment results. Ferrimagnetism and canted antiferromagnetism are the behaviors exhibited by minerals which are capa-

ALIGNMENT OF SPIN MOMENTS

1111 I l l IIIi Ill Ferromagnetism Net Moment

I

Ferrimagnetism

t

1111 lklk 1111 Antiferromagnetism

Canted Antiferromagnetism

Zero

Figure 3.2 Arrangement of spin moments associated with different classes of magnetic be-

havior.

3.2 Magnetic Minerals

29

ble of carrying magnetic remanence, the magnetic minerals. The ferrimagnetic minerals generally have high susceptibilities, high magnetization intensities, and relatively low coercivities, whereas the converse is true for antiferromagnetic minerals.

3.2 Magnetic Minerals a. Magnetite (Fe304) and the Titanomagnetites (xFe2TiO4.[1 -- x]Fe304) Magnetite is one of the end members of the ulvospinel-magnetite solid solution series (Fig. 3.3) and is the most common magnetic mineral on Earth. Magnetite is a cubic mineral with inverse spinel structure. The disordering (Curie) temperature varies more or less linearly with Ti content from -150~ for ulvospinel to 580~ for magnetite (Fig. 3.4a). Hence, ulvospinel is paramagnetic at room temperature. Titanomagnetites are ferrimagnetic at room temperature for values of x below about 0.8. Complete

Rutile TJ02

FeTi205 Ilmenite/ FeTi03~ Ulvospinel/

~

\

Oxidation to titanomaghemite series

/

Wustite FeO 3.3

~Pseudobrookite

-,%-,,,,.

\

Magnetite Fe304

Hematite Fe203

Maghemite

Figure Ternary diagram indicating the important iron oxide magnetic minerals, the titanomagnetite and titanohematite solid solution series, and the titanomaghemite oxidation trend.

30

3 Magnetization Processes and Magnetic Properties of Sediments

a 600

0.5

omagnetites

...,t~ 0_.4

400

c"

(b "_ . B

200 "~',

e-

~ E ,-

._o co

0 2.."

.. Sat..ura.t!on. , ' ~ ~ Magnetization (js, - ~ ~ ~ .

0"i" .....

!

0.0

0.2

Fe304

b E v -3

=-

|

0.6

.'% 0.8

Composition, x

0.50.4

-200 1.0 Fe2lqO4

600

Curie (Neel] Temp. (Tc)

0.3-

~ o

O

hematites

O

N

|

0.4

c-

400

B 200 "o t'b

-

,-,i

et~

E toO~

o e-

0.2. O.1.

Saturation /I Magnetization(Js~ /

60

t-

Antiferromagnetic ]

0 0.0

Fe203

012

014

O

-200

Ferrimagnetic o6

Composition, x

..

o8

o

-400

~o

FeTi03

Figure 3.4

Saturation magnetization (Js) and Curie temperature (To) as a function of (a) titanomagnetite composition and (b) titanohematite composition (after Nagata, 1961; Butler, 1992).

solid solution between ulvospinel and magnetite occurs only above 600~ and exsolution lamellae are common in igneous rocks. The exsolution is usually not between ulvospinel and magnetite, but between ilmenite and magnetite due to the high-temperature (deuteric) oxidation of ulvospinel to ilmenite, which usually accompanies exsolution (Haggerty, 1976a,b). As the ilmenite lamellae develop, the intervening titanomagnetite composition

3.2 Magnetic Minerals

3]

becomes progressively depleted in titanium. Further high-temperature oxidation results in the replacement of ilmenite and magnetite by aggregates of hematite, rutile, and pseudobrookite. High deuteric oxidation states are most common in acidic plutonic rocks which cool slowly and have high water content. A wide variety of deuteric oxidation states can be observed within a single lava flow and the controls are not clearly defined. In general, unexsolved titanomagnetites are most common in basic igneous rocks which have cooled rapidly, such as newly erupted submarine basalts. The exsolution process involves a change in grain size, grain shape, and composition of the titanomagnetite. The changes will generally involve a decrease in x, a decrease in grain size, and an elongation of the grain shape, and will therefore tend to increase the coercivity and blocking temperature of the grain. Therefore, the exsolution of magnetite and ilmenite by deuteric oxidation may serve to increase the stability of the remanence acquired during the process. The ilmenite separating the exsolved titanomagnetite rods is paramagnetic at room temperature and behaves as a nonmagnetic matrix for the ferrimagnetic titanomagnetite grains. Exsolution lamellae are a characteristic of titanomagnetite grains affected by deuteric oxidation. Low-temperature oxidation of titanomagnetites by weathering processes and/or oxidizing diagenetic conditions results in maghematization without exsolution, and subsequent formation of hematite. Titanomagnetites are a common constituent of igneous and metamorphic rocks and are therefore a prominent detrital component in sediments. Many sediments also contain low-titanium magnetite grains, which are uncommon in igneous rocks and are generally believed to be biogenic. Many organisms, including bacteria and molluscs, produce magnetite either by extracellular precipitation or as an integral part of their metabolism (e.g., Blakemore, 1975; Kirschvink and Lowenstam, 1979; Blakemore et al., 1985; Frankel, 1987; Lovley et al., 1987; Lovley, 1990; Bazylinski et al., 1988; Chang and Kirschvink, 1989; Sparks et al., 1990). Electron microscopy of magnetic extracts of lake sediments (Snowball, 1994) and a wide range of marine sediments (Vali et al., 1987; McNeill, 1990; Yamazaki et al., 1991) has shown that magnetite is often found in a restricted single domain (SD) grain size range, similar to the grain size of magnetite produced by living bacteria (Fig. 3.5). The grain size and low-titanium composition of this magnetite phase is such that it is an important carrier of stable magnetic remanence (see Moskowitz et al., 1988, 1994). o

b. Hematite (a-Fe203) and Titanohematite (xFeTiO3.[1 - x]Fe203) Mineral compositions intermediate between ilmenite and hematite are referred to as the titanohematites. Hematites have rhombohedral symmetry and corundum structure. As for the titanomagnetites, the disordering (N6el)

Marine Sedimentary Magnetite

Length (.m)

Length

t. 00 0.80 0.60

~.

0.40

SD

10000 8000 6000

MD

4000

~ , ~

0.20

oo oo

0.10

O. 08

0.06

~,~'~ : ~ ooo o o

+

+ +

,

2000

~tooo ~=~500

0.04 0.02

SP

o'2

o .4 k

]200

0 .'6

0 .'8

1.o

Axial Ratio

Modern Bacterial Magnetite

Length (~m)

. . . . . .

I. O0 O. 80 O. 60

Length toooo 8000 8000

MD

4000

0.40

0.20

= = 0~ = o

~'r

0.10 0.08

~

~ i"

=

oo ~

o

0.06

2000

~

,, d 1 ooo

" ~ 800 /~6oo

0.04 o. 02

SP 0

i

0.2

i

0.4

,

J200 A

0 6

0.8

1.0

Axial Ratio Figure 3.5 Comparison of size and shape of marine sedimentary magnetite and modern bacterial magnetite from electron microscopy (after Vali et al., 1987). SP: superparamagnetic; SD: single domain; TD: two domain; MD" multidomain.

3.2 Magnetic Minerals

~

temperature varies with x, from -218~ for ilmenite to 675~ for hematite (Fig. 3.4b). Compositions corresponding to values of x greater than about 0.8 are paramagnetic at room temperature. Other compositions are ferrimagnetic for values of x in the 0.5-0.8 range and antiferromagnetic for 0 < x < 0.5. Complete solid solution occurs only above about 1000~ and exsolution of the two end members is common except where x < 0.1 or x > 0.9. Titanohematites are associated with some metamorphic and plutonic rocks but almost pure hematite is much more abundant in metamorphic, sedimentary, and igneous rocks. Hematite is a very important as a remanence carrier in both clastic and chemical sediments. The hematite may occur in sediments as detrital specularite or as an authigenic pigment which grows during diagenesis. Hematites are antiferromagnetic, the weak net moment being due to slight canting of the spin moments (Fig. 3.2). The saturation magnetization is much lower than that of magnetite, and the remanent coercivity is significantly higher.

c. Maghemite (~'-Fe203) and Titanomaghemite (xFeTiO3.[1 - x]Fe203] Maghemite has the spinel structure of magnetite and the chemical composition of hematite. Maghemite and titanomaghemite are metastable phases, with maghemite and titanomaghemite inverting to hematite and magnetite, respectively, above 250~ The disordering (Curie) temperature, is therefore, difficult to determine but is believed to be about 640~ The magnetic properties are similar to those of magnetite, with slightly lower saturation magnetization and comparable remanence coercivity. Maghemite is common as a low-temperature oxidation product of magnetite in both igneous and sedimentary rocks.

d. Goethite (~-FeOOH) Goethite is the most common of the iron oxyhydroxides and is antiferromagnetic with a disordering (N6el) temperature of about 120~ (Hedley, 1971). Goethite has weak saturation magnetization and very high coercivity (Hedley, 1971; Rochette and Fillion, 1989). It dehydrates to hematite at about 300~ Goethite is an important product in low-temperature oxidation, particularly of iron sulfides, and is a constituent of manganese nodules and of the Pacific red clay facies, where it may precipitate directly from seawater. Lepidocrosite (y-FeOOH) is less common than goethite, has a N6el temperature of -196~ and is therefore not an important remanence carrier, but it dehydrates to maghemite and may therefore indirectly contribute to magnetic remanence.

34

3 Magnetization Processes and Magnetic Properties of Sediments

e. Pyrrhotite (FeSl+x, where x = 0-0.14) Fe7S8, the most magnetic pyrrhotite, is monoclinic and ferrimagnetic, with a Curie temperature of about 325~ Fe9S~0 is less important magnetically, being ferrimagnetic in a very restricted temperature range (--100-200~ The saturation magnetization of Fe7S8 is about ten times that of hematite and about half that of magnetite and maghemite (Clark, 1984). Remanence coercivity is generally greater than that of magnetite and maghemite (Dekkers, 1988, 1989). Pyrrhotite occurs in basic igneous rocks (Soffel, 1977, 1981) and can grow during sediment diagenesis in certain reducing environments (Kligfield and Channell, 1981; Freeman, 1986). The most common iron sulfide, pyrite (FeS2), is paramagnetic.

f. Greigite

(Fe3S4)

Greigite is ferrimagnetic and carries a stable magnetization in some lake sediments (Snowball and Thompson, 1990; Snowball, 1991) and in rapidly deposited clastic marine sediments (Tric et al., 1991a; Roberts and Turner, 1993; Reynolds et al., 1994). Coercivities are similar to those of magnetite and therefore generally somewhat less than for pyrrhotite. As for pyrrhotite, the Curie temperature is about 320~ and thermomagnetic curves are characterized by an increase in magnetic moment (and susceptibility) as the iron sulfide oxidizes to magnetite.

3.3 Magnetization Processes a. Thermal Remanent Magnetization (TRM) The remanent magnetization acquired by a grain during cooling through its blocking temperature is referred to as thermal remanent magnetization. The blocking temperature of a grain is the temperature above which the magnetic ordering in the grain is randomized by thermal energy during laboratory heating. The magnetic ordering of the grain will eventually become randomized at any temperature; the time that it is likely to take is referred to as the relaxation time (Tr), which is a function of coercive force (He), grain volume (v), spontaneous magnetization (Js), and temperature (T). Tr = To exp (vHcJs/2kT) where k is Boltzmann's constant and To is the frequency factor ('--10 -9 S). The exponential form of this equation means that, for a particular grain, the relaxation time increases logarithmically with decreasing temperature.

3.3 Magnetization Processes

35

The temperature at which the relaxation time becomes large relative to laboratory experimental time is referred to as the blocking temperature. For a rock or sediment sample, the grain population will have a blocking temperature spectrum reflecting the compositional and grain size range of the magnetic minerals. For weak magnetic fields, the TRM intensity is approximately proportional to the applied field, although lower cooling rates will enhance the TRM intensity. It is important to note that natural blocking temperature of a grain during igneous cooling will be lower than the laboratory-determined blocking temperature. The differences are greater at lower igneous cooling rates and become smaller when the blocking temperature is close to the Curie temperature (Dodson and McClellandBrown, 1980).

b. Chemical Remanent Magnetization (CRM) Chemical remanent magnetization is produced by an increase in grain volume at temperatures below the blocking temperature of the grain. The process is analogous to TRM acquisition but in the case of CRM the relaxation time [see Eq. (3.1)] is increased by change in grain volume (v) rather than ambient temperature (T). The logarithmic increase in relaxation time with grain volume leads to the concept of blocking volume, which is the critical grain volume at which the relaxation time becomes geologically significant. CRM is a very important acquisition process in sediments because magnetic minerals commonly grow authigenically during diagenesis. Magnetite, hematite, maghemite, goethite, and pyrrhotite can all grow authigenically in certain diagenetic conditions. Alteration during weathering can produce a CRM carried by goethite and/or hematite in a wide variety of rock types.

c. Detritai Remanent Magnetization (DRM) This process is an important source of primary magnetization in sediments. Detrital or biogenic magnetic mineral grains with a preexisting TRM or CRM can align themselves with the ambient field during deposition. The alignment occurs largely within the upper few centimeters of soft sediment rather than in the water column (Irving and Major, 1964). Magnetite, titanomagnetite, hematite, and maghemite are common carriers of DRM in sediments. The lock-in depth of DRM has been estimated by laboratory experiments (Kent, 1973; Lovlie, 1974; Tauxe and Kent, 1984), by comparing the position of reversals with oxygen isotope data in cores with varying sedimentation rates (de Menocal et al., 1990), and by observation of the damping of secular variation records (Hyodo, 1984; Lund and Keigwin, 1994). The estimates often lie in

~

3 Magnetization Processes and Magnetic Properties of Sediments

the 10-20 cm depth range below the sediment/water interface. The lock-in depth and width of the lock-in zone will depend on the grain size spectrum, sedimentation rate, and bioturbation depth.

d. Viscous Remanent Magnetization (VRM) VRM is an important source of secondary magnetization in rocks and sediments. Grains with relatively low relaxation time will tend to become remagnetized in the ambient weak field. The intensity of the VRM has been found both theoretically and experimentally to be proportional to the logarithm of time, and the proportionality constant (referred to as the viscosity coefficient) increases with temperature (Dunlop, 1983). Grains with volumes close to the blocking volume (i.e., close to the threshold between single domain and superparamagnetic behavior) are likely to be the most important carriers of VRM, although multidomain grains can also carry appreciable VRM. For grains with high coercivity, such as fine-grained hematite, the VRM acquired during the Brunhes chron can have high coercivity and blocking temperature up to about 350~ VRM components of the NRM will generally be aligned with the present Earth's field.

e. Isothermal Remanent Magnetization (IRM) This is a high-field magnetization process and therefore does not generally occur in nature except in the case of lightning strikes. As the properties of the IRM are useful for determining the magnetic mineralogy and grain size, IRM is produced in the laboratory using a pulse magnetizer, electromagnet, or superconducting magnet. The IRM is the remanence measured after removal of the sample from the magnetizing field. The intensity of the IRM will increase as the strength of the magnetizing field is increased until the saturation IRM (SIRM) is acquired. The magnetizing field required to attain the SIRM is a function of the remanent coercivity (Hcr) of sample grains.

f. Anhysteretic Remanent Magnetization (ARM) ARM is an artificial remanence produced in the laboratory by placing the sample in an alternating high field (typically 100 mT) with a low d.c. biasing field (typically 50-100 ~T) and allowing the alternating field to ramp down smoothly to zero. ARM is very useful because it has similar characteristics to TRM, and the characteristics of ARM can be used to determine grain size and composition of magnetic minerals present in the sample. The anhysteretic susceptibility (karm) is the ARM divided by the biasing field.

3.4 Magnetic Properties of Marine Sediments

37

g. Self-Reversal Mechanisms Self-reversal is a phenomenon where the remanent magnetization of a sample is antiparallel to the direction of the ambient magnetic field at the time of remanence acquisition. It is now clear that self-reversal is rare and the phenomenon is restricted to a few well-documented cases, particularly in titanohematites. Self-reversal can occur as a result of magnetostatic interaction between adjacent magnetic phases with different blocking temperatures and usually results from negative exchange coupling between a strongly magnetic, low Curie point phase and a weaker high Curie point phase with moment aligned with the ambient magnetizing field. The phases may be distinct mineral compositions or compositional gradients within grains. If the lower blocking temperature phase has greater volume and spontaneous magnetization than the higher blocking temperature phase, the net remanent magnetization can be antiparallel to the ambient field. The best documented cases of self-reversal are in a dacite containing titanohematite in the 0.45 < x < 0.60 compositional range (Uyeda, 1958) and in some oxidized titanomagnetite compositions (Schult, 1976; Heller and Petersen, 1982). Pyroclastics from the 1985 eruption of Nevado del Ruiz (Columbia) are reversely magnetized, and the self-reversal process has been replicated in laboratory-induced acquisition of TRM (Heller et al., 1986).

3.4 Magnetic Properties of Marine Sediments The magnetic properties of marine sediments depend not only on the nature of primary magnetic minerals contributed from detrital or biogenic sources, or by precipitation from seawater, but also on the diagenetic conditions which determine the alteration of primary magnetic phases as well as the authigenic growth of secondary magnetic minerals. The important remanence-carrying minerals in marine sediments are magnetite (Fe304), titanomagnetite, hematite (oz-Fe203), maghemite (y-Fe203), geothite (o~FeOOH), and iron sulfides such as pyrrhotite (FeS~+x) and greigite (Fe3S4). Of the magnetic minerals listed above, magnetite, titanomagnetite, greigite, and goethite can be primary remanence carriers in marine sediments. Hematite, pyrrhotite, and greigite occur as authigenic secondary minerals formed during early diagenesis, and goethite can be a product of low-temperature oxidation (weathering), particularly of iron sulfides. The magnetic properties of sediments from the four marine environments for which there is most information are reviewed below. a. Shallow Water Carbonates The majority of shallow water platform limestones are unreliable recorders of the geomagnetic field. However, there are several notable exceptions

~

3 Magnetization Processes and Magnetic Properties of Sediments

where early magnetizations have been recorded in this facies. For example, Silurian reef-slope limestones from Indiana yield magnetite magnetizations which predate compaction-related tilting (McCabe et al., 1983). Welldefined Plio-Pleistocene magnetostratigraphies have been acquired from shallow water limestones and dolomites recovered by drilling in the Bahama Bank (McNeill et al., 1988) and from the Mururoa Atoll (Aissaoui et al., 1990). Single domain magnetite of probable biogenic origin is thought to be the carrier of remanence in the partially dolomitized carbonates from the Bahamas, and the dolomitization process does not appear to remagnetize the carbonates (McNeill, 1990). A similar conclusion was drawn from study of Carboniferous limestones in Wyoming (Beske-Diehl and Shive, 1978). In the Mediterranean area, the "southern Tethyan" Mesozoic platform limestones generally carry very weak magnetizations, which are either too weak for precise measurement or are secondary and due to the weathering of pyrite to hematite and/or goethite. "Helvetic" limestones from the European Mesozoic continental margin carry a recent VRM with no resolvable primary magnetization components (Kligfield and Channell, 1981). The more marly shallow water Jurassic limestones from southern Germany carry a weak low-coercivity magnetization attributed to detrital magnetite; however, the purer limestone facies (bafflestones) carry a secondary magnetization due to both goethite and hematite produced by late oxidation of pyrite (Heller, 1978). A magnetite magnetization in marly shallow water Late Cretaceous limestones from the Munster Basin predates latest Cretaceous folding and therefore appears to be primary in origin (Heller and Channell, 1979). In North America, Paleozoic neritic and lagoonal limestone exposed in New Jersey and New York such as the Ordovician Trenton limestone (McElhinny and Opdyke, 1973; McCabe et al., 1983), the Siluro-Devonian Helderberg Series (Scotese et al., 1982), and the overlying Onodaga limestone (Kent, 1979) carry a synfolding or postfolding remanent magnetization with a direction implying remagnetization during Late CarboniferousEarly Permian time (Scotese et al., 1982; McCabe et al., 1983). These limestones contain spheroidal/botryoidal magnetite which grows after pyrite during diagenesis (Suk et al., 1990, 1993), and the particular diagenetic conditions which bring about this transformation have been associated with the migration of hydrocarbons during Alleghenian thrusting (McCabe et al., 1983). Comparison of the coercivity of IRM and ARM suggests that the magnetite is mainly in the MD size range (McElhinny and Opdyke, 1973; Kent, 1979). Curie point determinations and energy-dispersive X-ray analysis indicate that the carrier is a low-Ti magnetite (Scotese et al., 1982; McCabe et al., 1983). Remagnetized Paleozoic limestones from North

3.4 Magnetic Properties of Marine Sediments

~Q

America and Britain have characteristically "wasp-waisted" hysteresis loops (Fig. 3.6) (Jackson, 1990; McCabe and Channell, 1994). The origin of such wasp-waisted loops can be attributed to mixtures of grains with contrasting coercivities. In North American remagnetized limestones, the wasp-waisted characteristic of the loops is observed not only in the bulk rock but also in the "nonmagnetic" residue after magnetic extraction of the large (> few micron) sized spheroidal/botryoidal magnetite (Sun and Jackson, 1994). The large magnetite grains do not, therefore, appear to contribute to the characteristic wasp-waisted loop shape, which could be produced by mixing of SD magnetite with a volumetrically dominant SP magnetite phase (Jackson et al., 1993; McCabe and Channell, 1994; Channell and McCabe, 1994).

b. Pelagic Limestones and Calcareous/Siliceous Ooze Deep-sea (pelagic) limestones, and the unlithified analogues (calcareous ooze) recovered from ocean drilling, are the group of marine sediments that have been most thoroughly studied from a magnetic viewpoint. The interest in this sediment type stems from their efficiency as recorders of the ancient geomagnetic field and their usefulness as a means of establishing the correlation of polarity reversals to microfossil biozonations. There are many examples of the use of this type of sediment in magnetostratigraphic studies (e.g., Opdyke, 1972; Opdyke et al., 1974; Alvarez et al., 1977; Channell et al., 1979; Lowrie et al., 1982; Tauxe et al., 1983c). Many pelagic limestones and calcareous/siliceous oozes have rather simple magnetic properties, exhibiting a single magnetization component carried by ferrimagnetic magnetite or titanomagnetite, typically with hysteresis ratios lying in the pseudo-single domain (PSD) grain size range (Fig. 3.7). These sediments acquire their magnetization (DRM) by mechanical rotation of magnetite grains into line with the ambient geomagnetic field in the bioturbated upper few centimeters of soft sediment. As discussed above (w the depth (below the sediment/water interface) and size of the "lock-in zone" in which the grains lose their mobility will depend on grain size, magnetic mineralogy, sedimentation rate, and the extent of bioturbation. Estimates of this depth for pelagic sediments are generally a few tens of centimeters, which correspond to a few times 10 4 years, at typical pelagic sedimentation rates. Sediments which carry a magnetite DRM are often excellent recorders of the direction of the geomagnetic field at a time close to the time of deposition. Magnetite in pelagic sediments was thought to be entirely detrital (e.g., Lovlie, 1974); however, Curie temperatures of magnetic separates from deep-sea sediments often indicate the presence of almost pure magnetite

40

3 Magnetization Processes and Magnetic Properties of Sediments

Chatbum Limestone (UK) o

i

i

!

i

i

i

1'

1'

i

i

i

i

i

1

i

i

i"

i

:iiiiiiiiiiiiiiiiiiiiiiiiiii!iiiiiiiiiiiiiill 0

-2 -3 -300

-200

-100

0

100

200

300

H: Applied Field (mT)

60

Onondaga Limestone (NY)

40

20 E ""

0

~

-20 -40 -60 -300

-200

-100 0 100 H: Applied Field (roT)

200

300

Figure 3.6 Hysteresis loops from limestones remagnetizated during Late Carboniferous-Early

Permian time: top, Carboniferous limestones from Britain; bottom, Ordovician limestones from New York State (after McCabe and Channell, 1994).

3.4 Magnetic Properties of Marine Sediments

15

!

,

,

~

~

t

1

t

L,,.~

t

L

41

t

~

t

t

,

t

t

10".

.

S

Od

E

,r v

0

E

§

~E

'

-~o .

.

"................... ; ............................... . -

-6oo

-400

-200

0

Applied

200

Field

400

600

(mT),

0.6-1

O.S;

9

0.4gl

:E

u~

~r

PSD

~e

0.3'

,..1: %, %

-

o.z-

~o +

~ : @ I .AI

0

0.1-

MD 0-' 0

1

''I

2

. . . .

1 ....

3

4

'

' " " I '

S

''

'i

6

''

' ' i '

7

''

'1

8

Hcr/Hc Figure 3 . 7

(a) Typical hysteresis loop for white pelagic limestone (Maiolica Formation, Italy). (b) Hysteresis ratios for pelagic limestones of the Maiolica Formation, single domain (SD), pseudo-single domain, and multidomain (MD) magnetite fields according to Day et al. (1977) (data after Channell and McCabe, 1994).

42

3 Magnetization Processes and Magnetic Properties of Sediments

(Fe304). As pointed out by Henshaw and Merrill (1980), this is inconsistent with a totally detrital source as detrital source rocks (such as mid-ocean ridge basalts and igneous rocks exposed on land) contain titanomagnetite, rather than pure magnetite. The relatively recent discovery that several marine organisms, including bacteria and molluscs, precipitate nearly pure magnetite either extracellularly or as an integral part of their metabolism (Blakemore, 1975; Kirschvink and Lowenstam, 1979; Moskowitz et al., 1988, 1994) may provide the answer to this dilemma. Magnetite crystals of probably bacterial origin have been observed in pelagic sediments using transmission electron microscopy (Petersen et al., 1986; Vali et al., 1987; Yamazaki et al., 1991). These typically octahedral or parallelepiped-shaped magnetite grains are usually 0.05 to 0.1/~m across and within the single domain grain size range, which is optimal for retention of magnetic remanence (Fig. 3.5). Magnetite of bacterial origin may form at a few centimeters below the sediment-water interface at the transition from iron-oxidizing to ironreducing conditions (Karlin et al., 1987). It appears that nearly pure magnetite of biogenic origin is a major carrier of magnetic remanence in deepsea sediments. Peterson et al. (1986), distinguished two populations of magnetite grains in pelagic sediment from the South Atlantic: (1) a poorly sorted, irregular-shaped titanomagnetite of probable detrital origin and (2) a well-sorted, euhedral SD low-Ti magnetite of probable bacterial origin. Although the magnetites are the dominant remanence-carrying minerals in pelagic limestones and calcareous oozes, there is ample documentation that diagenetic alteration of primary magnetic minerals can occur in this environment. At typical pelagic sedimentation rates of the order of 10 m/ My, organic matter tends to get oxidized at the sediment/water interface. Hence, redox conditions in the soft sediment approach those in seawater, and under these conditions magnetite may be slowly but progressively oxidized. Tucker and Tauxe (1984) show a progressive oxidation of primary magnetite downcore at DSDP Leg 73 sites. Maghemite is the probable product of this oxidation, and the oxidation may occur by leaching of iron from the magnetic spinel phases rather than by incorporation of extra oxygen into the lattice. Maghemite has been adduced to occur in pelagic limestones cropping out on land (Channell et al., 1982a; Lowrie and Heller, 1982). The magnetic properties of Tethyan pelagic limestones exposed on land and in the Mediterranean area are strongly color dependent (Lowrie and Alvarez, 1975, 1977a; Channell et al., 1982a; Lowrie and Heller, 1982; Channell and McCabe, 1994). Red and pink pelagic limestones contain maghemite and hematite, in addition to the primary magnetite. The contrasting coecivities of magnetite and hematite produce wasp-waisted hysteresis loops (e.g., Fig. 3.1). White pelagic limestones are magnetite bearing

3.4 Magnetic Properties of Marine Sediments

43

(Fig. 3.7), do not appear to contain maghemite or hematite, but often contain optically visible pyrite, which readily weathers to goethite and finegrained hematite. The mineral goethite can also play a more primary role in the magnetization of pelagic sediments. It is a constituent of manganese nodules and can precipitate directly from seawater and, together with ironbearing clay minerals may be a precursor to pigmentary hematite (in the red limestones) and to pyrite (in the white limestones). Neither maghematization of primary magnetite nor authigenic growth of pigmentary hematite appears to occur in the white pelagic limestones. The remanence carried by pigmentary hematite in the red limestones appears, in some cases, to be due to growth of hematite in the uppermost few meters of soft sediment (Channell et al., 1982a). The dominant control on limestone color and on magnetic properties is probably organic content, the burial of organic matter being largely controlled by sedimentation rate.

c. Pacific "Red Clay" Facies Not all deep-sea sediments are efficient recorders of the geomagnetic field at the time of deposition. The lack of a primary magnetization may be due to the magnetic mineralogy or grain size of the detritus, to adverse diagenetic conditions which result in the alteration of primary magnetite, or to the growth of authigenic magnetic mineral phases. The so-called red clay facies occurs over a large part of the midlatitude Pacific (Davies and Gorsline, 1976). It is devoid of calcareous and siliceous microfossils and accumulated below the CCD at rates of about 25 cm/My. In this facies, the primary magnetization generally degrades at a few meters depth below the sedimentwater interface. The boundary between the primary and stable magnetic records often occurs in the later part of the Gauss chron and coincides closely with the late Pliocene onset of northern hemisphere glaciation (Opdyke and Foster, 1970; Kent and Lowrie, 1974; Prince et al., 1980). A number of different explanations have been offered to account for the degradation of the primary magnetization. Kent and Lowrie (1974) considered that oxidation of primary magnetite to maghemite accounts for the observed increase in VRM downcore. The SP-SD grain size threshold for maghemite is greater than that for magnetite, and therefore the oxidation process might be expected to increase the VRM contribution to remanence. In this scenario, the increased sedimentation rates in the upper part of the core characterized by stable primary remanence retard the oxidation (due to increased burial of organic matter). Johnson et al. (1975) also attributed the instability of the magnetization to the diagenetic growth of maghemite downcore, but considered the secondary magnetization to be a CRM rather than a VRM. Henshaw and Merrill (1980) have suggested that secondary

44

3 Magnetization Processes and Magnetic Properties of Sediments

ferromanganese oxides and oxyhydroxides [such as the jacobsite (MnFe204) solid solution series] are important authigenic magnetic phases and can carry a CRM which masks the primary DRM. Yamazaki and Katsura (1990) used a suspension method to determine the magnetic moment distribution and ferrimagnetic grain size in redeposited red clay. The mean grain diameter decreases below 1 m depth in the sediment, and this is consistent with changes in the frequency dependence of susceptibility, which indicates an increase in SP grains below this depth. These data imply that the increased importance of VRM downcore may be due to reduced grain size of eolian detrital magnetite in preglacial time (due to less vigorous atmospheric circulation), rather than to downcore maghematization.

d. Hemipelagic Sediments and Turbiditic "Flysch" Deposits In areas of high sedimentation rate and/or restricted bottom-water circulation, the burial of organic matter can result in reducing diagenetic conditions and the formation of iron sulfides. Paramagnetic pyrite is by far the most common iron sulfide in sediments, and it can grow by reduction of primary magnetite. The metastable iron sulfides mackinawite (FeS) and greigite (Fe3S4) are, unlike pyrite, capable of carrying magnetic remanence, as are some compositions of the more stable pyrrhotite (FeS~+~). Pyrrhotite coexists with magnetite and pyrite in young siliciclastic sediments from the Sea of Japan (Kobayashi and Nomura, 1972) and in Mesozoic "helvetic" limestones (Kligfield and Channell, 1981). Greigite in siliciclastic marine sediments from reducing depositional/diagenetic environments has been recognized in northern Italy (Tric et aL, 1991a), New Zealand (Roberts and Turner, 1993), and the North Slope of Alaska (Reynolds et al., 1994). Karlin and Levi (1985) have documented the progressive dissolution of primary magnetite, due to reducing diagenetic conditions, in hemipelagic silts and clays from the Oregon coast and in laminated diatomaceous oozes from the Gulf of California. Sedimentation rates were estimated to be 121 and 135 cm/ky, respectively. A decrease in NRM, SIRM, and ARM intensity occurs in the topmost meter as dissolution reduces the population of finegrained magnetite. Increases in coercivity and relative stability of ARM relative to IRM, from about 1 m depth in the cores to the base at 4 m, are interpreted in terms of a reduction in mean effective magnetite grain size as dissolution proceeds. A similar sequence of events has been observed at ODP Site 653, in a high sedimentation rate core from the Tyrrhenian Sea (Channell and Hawthorne, 1990). Turbiditic "flysch" deposits can carry a primary magnetization; however, the principal remanence carrier is usually multidomain (MD) titanomagnetite, hence the remanent coercivity is low and the sediments tend to

3.5 Magnetic Properties of Terrestrial Sediments

45

readily acquire VRM. In the Late Cretaceous Gurnigel Flysch, the VRM acquired in the Brunhes Chron dominates the NRM such that the primary magnetization cannot be resolved (Channell et al., 1979). In turbiditic sequences, it is generally good practice to sample the finest upper parts of individual flows; however, even in the finer silt horizons, MD titanomagnetite dominates the magnetite fraction. Turbiditic deposits are attractive for magnetostratigraphic study as their high sedimentation rates give high potential magnetostratigraphic resolution. The high sedimentation rates may, however, result in a high rate of burial of organic matter, sulfate reduction to sulfide, and the alteration of primary detrital titanomagnetite to iron sulfide minerals.

3.5 Magnetic Properties of Terrestrial Sediments a. L a k e S e d i m e n t s

In early studies of lake sediments, detrital hematite was considered to be the major carrier of remanence on the basis of the apparent observation of the Morin transition at -10~ (Creer et al., 1972). Later studies showed that the reduction in magnetization intensity at this temperature was due to the reorientation of grains by growth of ice crystals and that the dominant carrier is fine-grained single domain magnetite (Stober and Thompson, 1977; Turner and Thompson, 1979). The grain size of the magnetite ranges from multidomain to superparamagnetic in some North American lake sediments (King et al., 1982), detrital sources and bacterial activity being the controlling factors. Bacterial magnetite has also been documented in European lakes (e.g., Snowball, 1994). Sediments acquire a detrital remanent magnetization (DRM) by mechanical rotation of the magnetite grains in response to the ambient geomagnetic field. Although the DRM may be acquired within a few days in redeposited lake sediment (Barton et al., 1980), observations of magnetic coercivity of lake sediment from close to the sediment/water interface gave an estimate of lock-in time for a moderately stable DRM of about 100 years (Stober and Thompson, 1977). Lock-in times of a few hundred years may significantly distort the secular variation record particularly if the lockin process occurs over an appreciable depth range in the sediment. If the lock-in or moment fixing function can be estimated (by experiment) for a particular sediment, the secular variation record can be deconvolved by numerical methods (Hyodo, 1984; Lund and Keigwin, 1994). Comparison of the magnetic properties of natural and redeposited lake sediments has shown that the grains are probably not fixed in place by dewatering of the

~6

3 MagnetizationProcesses and Magnetic Properties of Sediments

sediment to a critical level, but possibly by gels that form in the pores of the wet sediment and are particularly prevalent in organic-rich sediment (Stober and Thompson, 1979). The process of sediment drying often disrupts the fidelity of the record; however, this may be due to alteration of remanence-carrying iron sulfides. Greigite is an important carrier of magnetic remanence in lake sediments (Giovanoli, 1979; Snowball and Thompson, 1988; Snowball, 1991). Greigite has coercivity similar to that of magnetite, however it can be recognized by thermomagnetic analysis and Mossbauer spectroscopy. Although greigite is an authigenic mineral in lake and marine sediments, it grows during early diagenesis and can carry a high fidelity record of the geomagnetic field (e.g., Tric et al., 1991b).

b. Loess Deposits The first magnetic stratigraphies of Chinese loess deposits resulted in drastic revision of the duration of Chinese loess deposition, from ---1.2 My to ---2.5 My (Heller and Liu, 1982, 1984). In addition, magnetic susceptibility records define lithologic loess/paleosol cycles which correlate with glacialinterglacial marine d180 record (Kukla et al., 1988; Liu et al., 1985). At first, it was thought that the susceptibility fluctuations were controlled by variations in the concentration of windblown detrital magnetite, due to increased loess deposition during glacials superimposed on a more constant background deposition of windblown magnetite (Kukla et al., 1988). It is now generally accepted that the enhanced susceptibility in the paleosols (during interglacials) is a result of authigenic "magnetic enhancement" as a result of elevated rainfall/temperature. The magnetic enhancement is generally attributed to bacterial production of fine-grained SP and SD magnetite during pedogenesis (Zhou et al., 1990; Heller et al., 1991; Maher and Thompson, 1991, 1992; Banerjee et al., 1993; Evans and Heller, 1994), although some authors consider maghemite to be important (Eyre and Shaw, 1994). The loess sediment has higher coercivities than intervening paleosols (Fig. 3.8) due to the greater contribution of hematite in the loess (Heller and Liu, 1984; Maher and Thompson, 1992).

c. Continental Red Beds Continental nonmarine clastic "red beds" often have stable magnetization carried by two distinct phases of hematite, one detrital and one authigenic (Collinson, 1965, 1974; Purucker et al., 1980; Tauxe et al., 1980). The magnetization of coarser grained (often drab-colored) sandstone units is often dominated by "specular" hematite, of presumed detrital origin, whereas finer grained units (often more reddened) tend to have CRMs dominated

3.5 Magnetic Properties of Terrestrial Sediments

,~ 4 0 -

47

Loess

Paleosol

OI

< 20E

-

._~ 0

0

,

~

-

'f. -20 I=

-40 -500

_

0 Field (re'r)

500

~0 . . . . . . . . . -500

0 Field (roT)

500

Figure 3.8 Hysteresis loops from Chinese loess and intervening paleosol (after Heller and Evans, 1995).

by fine-grained authigenic "pigmentary" hematite. Continental red beds have been important in magnetostratigraphic studies of the Paleozoic (DiVenere and Opdyke, 1990, 1991a,b), Mesozoic (e.g., Helsley and Steiner, 1974), and Cenozoic (Tauxe and Opdyke, 1982); however, the age of the hematite (and its magnetization) has been the subject of considerable debate (Purucker et al., 1980; Walker et al., 1981; E. E. Larson et al., 1982; Steiner, 1983; Lovlie et aL, 1984). Redeposition experiments of naturally disintegrated hematite-bearing sediments indicate that the hematite grains can carry a detrital remanence (DRM) closely aligned to the ambient field (Tauxe and Kent, 1984; Lovlie and Torsvik, 1984). Characteristic magnetization components in rip up clasts from Triassic and Carboniferous red beds are randomly directed, indicating an early origin for the magnetization (Molina-Garza et aL, 1991; Magnus and Opdyke, 1991). Probably the most convincing evidence for the penecontemporaneous acquisition of redbed remanence is the extensive studies of the red beds of the Siwalik system of Pakistan by Johnson et al. (1982, 1985), Opdyke et al. (1982), Tauxe and Opdyke (1982), and many others. The Siwalik sediments contain bright red hematite-bearing units which range in age from Middle Miocene (15 Ma) to Late Pliocene (2.5 Ma). The hematite magnetizations predate Pliocene-Pleistocene folding, and the Miocene sediments contain no evidence of detrital magnetite. Tauxe et al. (1980) carried out a conglomerate test and demonstrated an early magnetization in red siltstone clasts formed in an overbank environment and subsequently incorporated in the gray sandstone of a river channel. Upon thermal demagnetization, the sedimentary clasts were randomly, but stably magnetized, showing that the CRM directions were acquired early in the history of the sediment. Further evidence that the characteristic magnetization of these beds was acquired

~8

3 Magnetization Processes and Magnetic Properties of Sediments

early in their history is the fact that the magnetic stratigraphy from these sediments can be correlated to the GPTS derived from the oceanic magnetic anomaly record. Independent fission track dates on tufts in the sedimentary sequence confirm the correlation. Tauxe et al. (1980) suggest that the CRM of these sediments was acquired within 20 ky or less of deposition. This will not always be the case in reddened sediments, where authigenic pigmentary hematite can significantly postdate deposition.

l

Laboratory Techniques

4.1 Introduction Sampling in indurated rocks is usually conducted using a hand-held drill. A sample 5 to 10 cm in length is cored and left standing in the hole. The sample is then oriented using a slotted tube which fits over the core. Orientation devices usually have a plate on which a sun or magnetic compass is mounted and aligned with the slotted tube. The compass is leveled and the azimuth and dip of the sample are recorded. A mark is made on the sample along the slot, and the downward direction is marked on the sample. In volcanic areas, it is usually necessary to take sun compass readings, as local magnetic anomalies may distort the Earth's magnetic field. Most modern orientation devices are equipped for sun compass readings. The direction of a sundial shadow is recorded, along with the precise time of observation. The direction and dip of the bedding at the site are recorded to enable the bedding plane to be returned to the horizontal. Oriented hand-sized samples are often taken from outcrop and in this case the dip and strike of a flat surface are marked and recorded. The samples are then returned to the laboratory, where they are fashioned into a standard size, either cores with diameter 2.5 cm and length 2.3 cm or, alternatively, cubes of 8-15 cm 3. The direction and intensity of the natural remnant magnetization (NRM) are measured using a magnetometer. Most laboratories in North America, Europe, and Japan are equipped with cryogenic magnetometers in which the pick-up coils and SQUID sensors operate at liquid helium temperature (Goree and Fuller, 1976; Weeks et al., 1993). Other types of magnetometers such as the rock generator type (Johnson et al., 1948) or 49

50

4 Laboratory Techniques

fluxgate spinner magnetometers (Foster, 1966) are still used in some laboratories (see Collinson, 1983). The intensity of the sample magnetic field is measured in three orthogonal directions. This allows the calculation of the total moment of the sample as well as the declination and inclination of the magnetization relative to the orientation line on the core or cube. The magnetization direction of the sample is then rotated into geographic coordinates and rotated again, about the strike of the bedding, to account for bedding tilt. The NRM of rock and sediment samples is often the resultant of more than one magnetization component. Individual components may be acquired at different stages in the history of the sample. Components acquired during deposition of the sediment or cooling of the igneous body are referred to as "primary." Those associated with later geological events such as sediment diagenesis, deformation, uplift, and weathering are referred to as "secondary." The objective of this chapter is to provide a practical guide to methods of determining (1) the component content of a magnetization and (2) magnetic minerals which carry the magnetization components. For detailed discussions of mineral magnetic properties and of instrumentation, see Collinson (1983), O'Reilly (1984), and Dunlop (1990). Information concerning modes of occurrence of magnetic minerals in marine sediments has increased considerably in the last few years (Ch. 3) such that an insight into the magnetic mineralogy is an important step in determining the relative age of resolved magnetization components. In magnetostratigraphic studies, it is necessary to establish which, if any, of the resolved magnetization components are primary and can be associated with the depositional or cooling age of the sedimentary or igneous rocks. Estimating the age of magnetization components is a critical step in paleomagnetic and magnetostratigraphic studies. Determination of the mineralogy of magnetic carriers does not generally provide unequivocal evidence for (but is often a guide to) the relative age of magnetization components. Field tests (the tilt test and conglomerate test) (w 5.3) are additional important means of constraining the age of magnetization components. Note that the normal and reverse directions recorded by a particular component do not necessarily indicate that the component is primary; on the other hand, the absence of antiparallel reverse directions may be an indication that the resolved components are secondary, or influenced by secondary components.

4.2 ResolvingMagnetization Components a. Alternating Field Demagnetization The strength of the external magnetizing field which is required to reduce the remanent magnetization of a population of grains to zero is referred

4.2 Resolving Magnetization Components

5|

to as the remanent coercivity, or the coercivity of remanence (w 3.1). The remanence direction of individual grains will be altered by different external field strengths depending on the "coercivity" of the individual grain, which is controlled by composition and grain size. A natural sample will have a "coercivity spectrum" reflecting the range of magnetic grain sizes and compositions present in the sample. This dependence of coercivity on composition and grain size is utilized in alternating field demagnetization to resolve the magnetization components carried by different grain populations. If the NRM of the sample is a composite of individual magnetization components which have distinct coercivity spectra, the alternating field method will be suitable for resolving the individual components. The normal procedure is to progressively demagnetize the sample by a stepwise increase in peak alternating field--the peak field is smoothly ramped down to z e r o ~ i n field-free space in order to avoid acquisition of ARM (Fig. 4.1), and the remaining remanence measured after each step. In a natural sample, grain remanences with coercivity lower than the peak alternating field will be reoriented as the peak field is ramped down, and their contributions to the remaining remanence will be randomized (Fig. 4.1). The net effect is to "demagnetize" those grains. As the peak demagnetizing field is progressively raised, a larger proportion of the total magnetic grain population becomes demagnetized and ceases to contribute to the remanence of the sample. There are two principal types of alternating field demagnetizers in current use. For both types, a Mu metal shielded solenoid generates an alternating field, with a frequency of a few hundred hertz, which is smoothly ramped down to zero from a designated peak alternating field. The sample is either placed in this field in three steps, such that the alternating field acts along the three orthogonal axes of the sample, or rotated in a specially designed tumbler about three (or in some cases two) orthogonal axes simultaneously, within the solenoid. The rotation of the sample within the coil system has the advantage of reducing the tendency for ARM acquisition; however, the observation that spurious remanent magnetization can be acquired along the axis about which the sample is rotated (Wilson and Lomax, 1972) has led to wide use of demagnetizers in which the sample is not rotated and is demagnetized in three individual steps along the orthogonal axes of the sample. The availability of highly efficient Mu metal shields has tended to negate the advantage of rotational methods in reducing the problem of ARM acquisition. For a full description of the design of alternating field demagnetizers, see Collinson (1983). b. Thermal D e m a g n e t i z a t i o n The time required for magnetic remanence of a grain to decay in zero external field (the relaxation time) is a function of ambient temperature,

52

4 Laboratory Techniques

~

i~ ~gran i~~

Peak alternating field

-H

+H

~

I

"'I'-'--~-----i

|

~

.______t_.__~~

I

Last ti me "-~=:::::~Z~Z grain A is I affected I , ~ I t

I

>

ta m

I-

Figure 4.1 Decay of alternating field strength during AF demagnetization. As the experiment is conducted in Mu metal shields in the absence of the Earth's field, the net magnetic moment of a large population of grains with coercivities less than the peak field is zero (after Van der Voo, 1993).

grain size, and composition (Fig. 4.2). The blocking temperature (or unblocking temperature) is the temperature at which the relaxation time of the grain falls within the time range of laboratory experiments (---103104 s). As the function is exponential, the exact length of time of the experiment does not, within reason, critically affect the observed blocking temperature. Two populations of magnetic grains with different origin, and therefore different composition and/or grain size, might be expected to have distinct blocking temperature spectra. This principle is utilized in thermal demagnetization to separate magnetization components carried by different magnetic grain populations. Normal practice is to heat the sample in progressive temperature increments (until the magnetization falls below the magnetometer noise level), allowing the sample to cool in field-free

4.2 Resolving Magnetization Components

MAGNETITE

1O0

i

200

i

|

300

|

.!

10my

,0~,~ ~ ~ ~

l\ \ \ i~ 1too

v

|

100

XX !

w

i

200

i

400

i

!

500 |

lby

105y

103y

1~ -lh

-10s 300

v

400

HEMATITE

200

9

10my

300

400

9

r

500

Blocking temperature (~

100

53

500 i

I

600 9

10my 105y t-,.

-10my

X

105y

103y

103y

~r9 10y \ .Ot~ Imo "

J~

1

1

lh \

10$

1 100

I

i

200

9

l

300

!

9

400

,

,

500

Blocking temperature (~

/

10y

-Imo -lh

k -lOs

600

Figure 4.2 Blocking/unblocking contours joining time-temperature points for populations of SD magnetite and hematite (after Pullaiah et al., 1975) illustrating that high temperatures in the laboratory time scale (say 1 hour) can unblock a thermoviscous magnetization acquired at lower temperatures on geologic time scales (say 10 My).

space after each heating increment. The remanence of all grains with blocking temperatures less than the maximum temperature will become randomized as the sample cools in field-free space. The magnetization measured after each heating step can be used to determine the directions of magnetization components with distinct blocking temperature spectra.

~

4 LaboratoryTechniques

As for alternating field demagnetizers, the advent of highly efficient Mu metal magnetic shielding has greatly improved and simplified the shielding of thermal demagnetization equipment. Most thermal demagnetizers are built inside a long (~2 m) Mu metal shield and a cylindrical furnace cap~ible of holding the samples lies adjacent to a cylindrical blown-air cooling chamber. The sample boat is pushed from the furnace into the cooling chamber inside the Mu metal shielding, and samples can be cooled while another set of samples is being heated. The furnace must be noninductively wound and the ambient field in the cooling chamber should not exceed a few nT, or a partial TRM may be acquired by the samples.

c. Chemical Demagnetization Leaching with hydrochloric acid can provide a means of selective destruction of magnetic grains. The method is applicable to siliceous rocks, particularly red sandstones, which contain a secondary magnetization due to authigenic growth of fine-grained, pigmentary hematite. The dissolution rate of coarse-grained hematite (specularite) of detrital origin is lower than that of the finer-grained, pigmentary hematite of diagenetic origin, due to the higher ratio of surface area to grain volume for finer grains. Leaching in 3N to 10N hydrochloric acid, and remanence measurement at certain time intervals during the leaching, provides a means of resolving components with distinct "leaching spectra." Slots are cut in the sample to maximize the surface area of the sample in contact with the acid.

d. Orthogonal Projections The most widely used method for analyzing changes in magnetization direction and intensity during demagnetization (using any of the above methods) is the orthogonal projection (Wilson, 1961), also referred to as the Zijderveld diagram (Zijderveld, 1967). This is a simple but extremely useful means of displaying the change in magnitude and direction of the magnetization vector during demagnetization. The end points of the magnetization vector are plotted both on the horizontal plane, where the axes are N, S, E, and W, and on the vertical plane where the axes are Up and Down (Fig. 4.3). Straight-line segments on these plots indicate that the magnetization vector being removed in this interval has constant direction. This direction can be computed and defined as a component of the magnetization. It is important to note that straight-line segments on orthogonal projections may not always represent single components. They may also represent the composites of two or more components with identical demagnetization (coercivity, blocking, or leaching) spectra. Invariably, in a particular projection, there

4.2 Resolving Magnetization Components

55

N dJ/dHor dJ/dT

~

a

Up, E

~'~. A

ze9 3 / J

2

N

Comp B O t 23 456 789 Peak alternating field (H) or temperature (T)

i

B "6 9

N

dJ/dHor dJ/dT

~

..'~

b

~~. 9

0 , Z 3 4 5-6 r 8 9 Peak alternating field (H) or temperature (T)

"

2

N

B 5"~/6 6 -y

N

dJ/dHor dJ/dT

1

Up, E

-

!

Up, E

2

C

,/4

0 m 2 ] 4 5 6 r 8 9 Peak alternating field (H) or temperature (T)

+

r

6

B e' / i

Figure 4.3 Equal area and orthogonal projection of demagnetization data for (a) nonoverlapping, (b) partially overlapping, and (c) overlapping coercivity or blocking temperature spectra (after Dunlop, 1979).

may be several straight-line segments indicating the presence of multiple components. Curvature of the lines at the junction between two straightline segments indicates overlap in demagnetization spectra (Fig. 4.3). The magnetization components cannot be resolved if the overlap in demagneti-

~6

4 Laboratory Techniques

zation spectra is too great. If the blocking temperature spectra overlap, this does not necessarily mean that the coercivity spectra will overlap, and the choice of demagnetization technique will be critical to whether the component content of the magnetization can be resolved. The points on the orthogonal projection which define straight-line segments are commonly picked by eye. The component direction associated with a particular straight-line segment can then be computed by a "leastsquares fitting" technique (e.g., Kirschvink, 1980). This method gives a maximum angular deviation (MAD) value which provides an estimate of the goodness of fit to the straight line to the points on the projection. MAD values greater than ---15~ usually indicate that the component is poorly defined.

e. Hoffman-Day Method In the special case in which three magnetization components with overlapping demagnetization spectra are present, the Hoffman-Day method (Hoffman and Day, 1978) may help to resolve the component directions. The method involves plotting subtracted vectors on an equal area stereographic projection, at each stage in the progressive demagnetization process (Fig. 4.4). In a demagnetization interval characterized by two overlapping demagnetization spectra, the subtracted vectors will plot on a great circle between the two component directions. If three overlapping demagnetization spectra are present (Fig. 4.4), the intersection of the two great circles can give an estimate of the direction associated with the component (B) with intermediate blocking temperature, even though this component direction is not directly resolved by the demagnetization process. Orthogonal projection in this case would not resolve the intermediate component.

f. Remagnetization Circles If two magnetization components with partially overlapping demagnetization spectra are present in the sample, the resultant vector will trace a great circle in stereographic projection as the overlapping interval is demagnetized. If one of the two components has dispersed directions, the great circles from individual samples will converge to give the direction of the undispersed component (Fig. 4.5). The method is applicable in the special case where one of the components is dispersed. If both components have uniform directions, there will be no appreciable convergence of the great circles, and the method cannot be used on its own. The dispersion of one of the components can occur in a number of situations. For example, the

4.2 Resolving Magnetization Components

Components

C

B

57

A

222

Coercivity or blocking temperature spectra Components

C BA

Coercivity or blocking temperature spectra Components

CBA

Coercivity or blocking temperature spectra

Figure 4.4 Difference vector paths corresponding to (a) nonoverlapping, (b) partially overlapping, and (c) overlapping coercivity or blocking temperature spectra (after Hoffman and Day, 1978).

primary component in conglomerate clasts will be dispersed relative to the secondary component, which postdates conglomerate deposition. In this case the secondary component could be resolved by remagnetization circle analysis. The method can also be applied in circumstances in which one of the two components postdates and the other predates the folding. In theory, both components can be resolved by applying the method both before and after tilt correction. Remagnetization circle analysis is most often used in conjunction with individual component directions resolved from orthogonal projections. If curvature of lines on orthogonal projections precludes the resolution of individual components for some but not all samples from a site, remagnetization circle analysis of these direction trends can augment and improve the estimate of the site mean direction derived from the orthogonal projections, using the procedure of McFadden and McE1hinny (1988).

58

4 Laboratory Techniques

270 ~

90 ~

','~, ,,

i!!',,','%,

,,

180 ~

Figure 4 . 5 Remagnetization circles obtained by AF demagnetization of a two-component magnetization in which one component is dispersed. Dashed segments are great circle extrapolations of the data (after Halls, 1976).

4.3 Statistics The techniques outlined above are designed to isolate individual components of magnetization. An individual component is often combined with others from the same horizon or from a particular sampling site. These vectors are represented as unit vectors and often displayed as points on an equal area (stereographic) projection, representing points on a sphere. Fisher (1953) developed a method for analyzing the distribution of points on a sphere. He predicted that these points would be distributed with the probability density P given by K

P = 4rr sinh(K) e Kc~

(4.1)

where ~ is the angle between the direction of magnetization of a sample and the direction (qJ = 0) at which the density of the observations is a maximum. If the density of the points is axially symmetrical about the true mean direction, with the points being distributed through 360 ~, then the probability of finding a point between 0 and 0 + 60 is:

4.3 S t a t i s t i c s

59

Po 60 = (K/2 sinh(K))e Kc~176 sin 0 60.

(4.2)

The parameter K determines the distribution of the points and is called the "precision parameter." If K = 0 then the points are distributed evenly over the sphere and the directions are random. If the values of K are large, the points are clustered tightly about the true mean direction. If a group of points representing directions of magnetization are distributed on a sphere, the best estimate of the mean direction of magnetization (the center of these points) can be found by the vector addition of all directions. The direction of magnetization of a rock is given in terms of the declination D measured over 360 ~ clockwise from true north and the inclination I measured positive (negative) downward (upward) from the horizontal. The unit vector representing this direction can be determined by its three direction cosines: north component

l = cos D cos I

east component

m = sinDcosI

down (up) component

n = sin I.

The resultant direction of N such directions may be determined by summing the direction cosines and is given by: N

li

X = 1/R Z i=1 N

Y = 1/R Z

(4.3)

mi

i=1 N

Z = 1/R Z

ni.

i=1

The vector sum will have a length R which is equal to or less than the number of observation and is given by: R2 =

li

t 2+

mi

t2+

i=1

ni

)2.

(4.4)

=

The declination and inclination of the vector mean direction R are given by: N

Tan DR = Z

N

m,IZ - ,

i=1

(4.5)

i=1 N

Sin IR = 1/R 2

ni.

i=1

The best estimate k of the precision parameter K for k > 3 is:

(4.6)

60

4 Laboratory Techniques

k = (N-

1 ) / ( N - R).

(4.7)

The angle from the true mean within which 95% of the directions will lie is given approximately by: qJ95 = 140/(X/k) degrees

(4.8)

This value represents the angle from the mean beyond which only 1/20 of the directions will lie. Fisher (1953) has shown that the true mean direction of a group of (N) observations will lie within a circular cone about the resultant vector R with a semiangle a for a certain probability level (1 - P) when k > 3. P is usually taken as 0.05 and if a is small, the circle of 95% confidence is given by: c~95 = 140/(Vk--N).

(4.9)

The value of Og95is often used to determine whether two sets of directions differ from one another. If the Ct95values of two mean directions of magnetization do not overlap the mean directions of magnetization are usually believed to be significantly different at the 95% confidence level. A more rigorous test of whether two populations are different was proposed by Watson (1956a). Fisher statistics are almost universally used in paleomagnetic analysis; however, the Fisher distribution is not ideally suited to paleomagnetic data for two reasons. Firstly, paleomagnetic directions are not always symmetrically distributed about the mean direction; elliptical distributed data sets are commonly observed. Secondly, paleomagnetic directions are often bimodal comprising both normal and reverse polarity directions. The Bingham distribution (Bingham, 1974; Onstott, 1980) is better suited to paleomagnetic data as noncircular (elliptical) and bimodal data sets are allowed. In this treatment, the "orientation matrix" (computed as the normalized matrix of sums of squares and products) is similar to the moment of inertia matrix with unit mass assigned to each data point on the sphere. The largest (principal) eigenvector provides a very satisfactory estimate of the mean direction for the population; however, the estimation of confidence limits about the mean is more problematic, resulting in the low popularity of this method in paleomagnetic analyses. Tauxe et al. (1991) advocated the use of "bootstrap" methods to calculate the confidence limits of paleomagnetic mean direction. The advantage of the method is that less stringent assumptions are made concerning the nature of the distribution (the probability density function). The disadvantage is that it is not suitable for data sets smaller than "a few dozen." The method is therefore inapplicable in cases where less than --30 magnetization components are resolved (samples collected) at a paleomagnetic sampling site. The bootstrap method involves random sampling of the data set (population: n) to produce a large number (~n 2) of pseudo-data sets each with

4.4 Practical Guide to the Identification of Magnetic Minerals

6|

the same population (n) as the original data set. Obviously, some data points may occur more than once in a given pseudo-data set. The parameter of interest, such as the mean direction, is calculated for each pseudo-data set, and the variability is used to calculate the confidence limits. For each pseudo-data set, the normal and reverse modes are separated about the normal to the principal eigenvector. The mean direction for each mode in the pseudo-data set is estimated as the sum of unit vectors, as in the Fisher procedure [Eq. (4.3)]. The mean of the bootstrapped means, the principal eigenvector of the orientation matrix derived from the bootstrap means, and the mean of the original data set should be very nearly the same. The confidence circle (a95) associated with the mean is then estimated from the bootstrap means assuming that they have a Kent (1982) distribution, which is the elliptical analogue of the Fisher distribution.

4.4 Practical Guide to the Identification of Magnetic Minerals Because magnetic mineral phases tend to have characteristic modes of occurrence, the identification of magnetic remanence carriers is a useful clue to the age of magnetization components. The following is designed to provide a guide to the identification of common remanence-carrying magnetic minerals. The focus is on convenient methods which are useful in sediments where the magnetic mineral grains are well dispersed and in low concentration. Most of the methods utilize instruments which are commonly available in paleomagnetic laboratories. In weakly magnetized sediments, the characteristics of artificially imposed remanences (such as ARM and IRM) are very useful for determining the magnetic mineralogy of the sample. However, in order to determine the carriers of NRM components, the characteristics of the NRM itself must be determined as fully as possible.

a. N R M Analysis As described above, progressive demagnetization is the standard procedure for determining the component content of the NRM. Thermal and alternating field demagnetization characteristics indicate the blocking temperature and coercivity spectra, respectively, of the NRM components. Blocking temperature is a function of composition and grain size of the magnetic carrier. For SD grains of constant composition, the blocking temperature will be proportional to grain size [Eq. (3.1)]. Due to the large grain size of hematite at the SD/MD threshold (---2 mm), the SD range for hematite covers most of the remanence-carrying grain sizes encountered in sediments. Therefore, in the case of hematite, blocking temperature is a useful guide to grain size. For titanomagnetites (xFezTiO4.(1-x)Fe304) the block-

6~

4 Laboratory Techniques

ing temperature is a function of composition and grain size. As x decreases, Curie point increases almost linearly (Fig. 3.4), and as blocking temperature cannot exceed the Curie point, the available blocking temperature range will increase as x decreases. For SD grains of uniform shape and composition, the blocking temperature will be a function of grain shape as well as grain volume. In the MD range, the rapid decrease in coercivity with increasing grain size might be expected to give rise to an accompanying decrease in blocking temperature. The blocking temperature range in itself is not generally diagnostic of magnetic mineralogy, other than the fact that the maximum blocking temperature cannot exceed the Curie or N6el point of the mineral composition. In the case of minerals with relatively low Curie or N6el points, such as geothite (-120~ the blocking temperature range is rather diagnostic because it is so restricted. Mineralogical changes as a result of heating during thermal demagnetization can obscure the blocking temperature spectrum. For example, maghemite alters to hematite at about 300~ iron sulfides will oxidize in air to form magnetite at about 300~ and hematite tends to convert to magnetite at temperatures above 600~ These phase changes, if identified as such, can be useful as a means of identification of magnetic minerals. The coercivities of remanence of hematite and geothite are typically about two and three orders of magnitude greater, respectively, than that of SD magnetite. Coercivity of remanence in naturally occurring hematite and titanomagnetite is largely a function of grain volume in the case of hematite and grain shape and volume in the case of titanomagnetites. For magnetites, coercive force (He) increases with x and with decreasing grain size in a wide MD grain size range (0.1-1000/xm) where Hc is proportional to d -~ with crushed grains having higher coercivity than hydrothermally grown or naturally occurring crystals (Heider et al., 1987) (Fig. 4.6). Below about 0.05/zm, the coercivities will decrease sharply close to the SP-SD grain size threshold (see Maher, 1988). Although demagnetization of remanence carriers from AF or thermal demagnetization data is rarely unequivocal, combinations of these data from the same samples can be fairly diagnostic. For example, in Fig. 4.7, three components, A, B, and C, are apparent in the orthogonal projection of thermal demagnetization data. The blocking temperature spectra of components B and C overlap. In the AF demagnetization diagram for a specimen from the same sample, only components A and C are apparently resolved. Component B has low blocking temperature (200-300~ and high coercivity (>80 mT). In view of the discussion above, fine-grained hematite is a possible carrier for this component. This is consistent with the red pigment of the limestone sample and suggests that this component is a secondary component of diagenetic origin. Component C has a maximum

4.4 Practical Guide to the Identification of Magnetic Minerals

50-

,t

20 = - - ' " k

o.

10 - Precipi s - and annealed "-"

: "...

O

1

._

0.s

0

0.2

I.--

O

MAGNETITE

-~ -,-.i., - -

-- -.

2

I.--

63

Crushed

~.. J - .-,-- X

grains

~

Hydrothermally recrystallized

~ x

X

x x

0.1 0.05

~

\ x Ix

Natural single crystals 9

i

,Jt~tJt|

Jt

,

,,,,,.,I

I

o.+

+

i

~,,,,.I

.,i.

+o

,

, ,,ii~|

~ i L .....

~oo

I

lOOO

Grain diameter (microns)

Figure 4.6 Coercive force as a function of grain diameter for magnetite (after Heider et aL, 1987). The apparent higher coercivities of crushed grains are attributed to their stressed state and adhering finer fragments.

N, Up

Thermal demagnetization

2"~oi4 ~ W

W

"

Component B / ' 2

0

0

~

~

~ e

Alternating field demagnetization

~

N Up

1"1o22......~

./'S~'n

t

C - 1 "10-4

~

ComponentC

S, Down

A/m

-2"10

-4

Horiz Vert

9= Component A 9A ComponentB 9o Component C

Component~A S, Down

Figure 4.70rthogonal projection of a three-component magnetization (from a pink pelagic limestone) illustrating contrasting response to thermal (temperatures in ~ and alternating field (peak fields in mT) demagnetization (after Channell et aL, 1982a).

64

4 LaboratoryTechniques

blocking temperature of over 600~ and is therefore carried, at least in part, by relatively coarse-grained hematite. The coercivity spectrum of component C suggests that a low-coercivity mineral (magnetite) contributes to this component. Component A has low blocking temperature and low coercivity, and the in situ direction is close to that of the present field; therefore the component is a recent (probably viscous) overprint. It is important to note that the more primary magnetization component is not necessarily the final component revealed by the demagnetization procedure. In the example mentioned above, the secondary component C has the highest coercivity of the components present in the sample.

b. Js/T Curves The change in saturation magnetization with temperature is an important indicator of magnetic mineralogy, because it is independent of grain size and shape. The temperature at which the saturation magnetization disappears is the Curie or N6el temperature, which is the temperature at which thermal energy precludes spin moment alignment. This temperature is characteristic of the magnetic mineralogy. Most Curie balance systems are not sufficiently sensitive to determine Curie or N6el points from weakly magnetized sediments without extraction and concentration of magnetic minerals. Magnetic extraction procedures are inevitably grain size and compositionally selective, and therefore the Js/T curves may not be representative of the magnetic mineral content of the sample. Significant improvements in Curie balance sensitivity have, however, been achieved by sinusoidally cycling the applied field and making continuous drift corrections through Fourier analysis of the output signal (Mullender et al., 1993). Horizontal balance systems are generally preferred to vertical systems, as they reduce weight loss problems arising from dehydration of clay minerals. For titanomagnetites and titanohematites, Curie and NEel temperatures indicate the composition (x) (Fig. 3.4), and in favorable circumstances Curie temperatures corresponding to several magnetic phases can be observed (Fig. 4.8). Alteration of magnetic minerals during heating may obscure the Curie or N6el point of the original mineral and produces irreversible curves. This irreversibility can be useful for detecting the presence of maghemite, which converts to hematite at about 300~ (Fig. 4.8), and for determining the presence of iron sulfides, which alter to magnetite when heated in air. Ideally, irreversible curves should be rerun in vacuo or in an inert gas atmosphere.

c. IRM Acquisition and Thermal Demagnetization of 3-Axis IRM The shape of the curve of acquisition of IRM with increasing magnetizing field is a measure of the coercivity spectrum of magnetic minerals present

4.4 Practical Guide to the Identification of Magnetic Minerals

[]-magnetite "',~.exsolvin~heatginon g

Js

,

' ' ' 5~0' 1=_ I

Maghemite ' ~ ~ar~igtO 5O0

andhematite L ~

~

" ~ PlYrrrhO~i tteo Js ~~tite

Js [ ' ~

65

500 k

" 5~)o' T'oc

hematiteUnSaturated ,

~

,

500 !

v

ToC ,

Paramagnetic 5O0

500

Toc

Figure 4.8 Typical saturation magnetization-temperature (Js/T) curves for various minerological compositions (after Piper, 1987).

in the sample (Dunlop, 1972). The common magnetic phases present in sediments, SD magnetite, MD magnetite, fine SD hematite, coarse SD hematite, and goethite have rather characteristic and distinct coercivity spectra and IRM acquisition curves (Fig. 4.9). The intensity of the SIRM and the magnetizing field at which the SIRM is acquired vary with mineralogy and grain size. Thermal demagnetization of a composite 3-axis IRM (Lowrie, 1990) is very useful for determining magnetic mineralogy. An IRM is applied sequentially along three orthogonal axes of the sample using decreasing magnetizing fields, typically 1 T, 0.4 T, and 0.12 T. Thermal demagnetization of the three orthogonal "soft," "medium," and "hard" IRMs can give insights into the magnetic mineralogy (Fig. 4.10).

d. Hysteresis Loops The advent of the alternating gradient force magnetometer 0xMAG) allows hysteresis loops to be generated for small (---50 mg) samples, even for weakly magnetized sediments. Saturation magnetization (M~), saturation remanence (Mrs), and coercive force (Hc) can be determined from the loops, and back field experiments on the/xMAG can yield values of remanent coercivity (Hcr) (Fig. 4.11). For samples containing magnetite only, hystere-

66

4 Laboratory Techniques

JR'(

JR(A/m)

0.10

aJ/aH

I00

(%)

Magnetite and hematite

0.05

1

2

3

~J/,~H (%) -IOO

0.15 -

50

0

/

0"30t

5O

1

40-~H(T)

=

3

,

O~H

(1")

JR(A/m)

0.201 JR(A/m)

L

/

,oo

HI%)

0.30

(%)

_

0.1

0

1

2

3

0-~H-- (1)

0

" . . .1. . . 2 "' 3'

e

i0

~4

)~-H (T) 5

Figure 4.9 Acquisitionof isothermal remanent magnetization (IRM) for various mineralogical compositions in limestones (after Heller, 1978).

sis parameters determined in this fashion can be plotted on a hysteresis ratio plot to estimate the domain state of the magnetite population (Day et al., 1977) (Fig. 4.12). e. A F

Demagnetization of IRM and

ARM

A modification of the Lowrie-Fuller (1971) test introduced by H. P. Johnson et al. (1975), provides a useful means of determining the relative importance

of SD and MD grains in magnetite-bearing rocks and sediments. The original Lowrie-Fuller test is based on the observation and theory that the

4.4 Practical Guide to the Identification of Magnetic Minerals

67

a f 150"

E < E

100

t_

~

50

1

2

3

4

5

Magnetizing Field (T)

b

~ ~=~

100

=

Soft (0.12 T) Medium (0.5 T) Hard (1.0 T)

,,.=,,

<

50 L_

9

0

-

,

'

v

200 400 600 Demagnetizing temperature (~

Figure 4.10 For a limestone sample containing pyrrhotite and hematite, (a) acquisition of isothermal remanent magnetization (IRM) and (b) thermal demagnetization of a 3-axis IRM (after Lowrie, 1990).

stability to AF demagnetization of thermal remanent magnetization (TRM) in SD magnetite increases with decreasing magnetizing field, whereas the converse is true for MD grains. This observation is not strictly a domain state delimiter; however, it is a useful grain size parameter (Dunlop, 1986). In the modified version of this test (H. P. Johnson et al. 1975), it is assumed that the stability of strong field TRM will be similar to that of SIRM and that the stability of low-field TRM will be similar to that of ARM. The modified test compares the stability of AF demagnetization of SIRM and ARM, is very simple to apply, and avoids the need to heat the sample, which may produce mineralogical change. According to the H. P. Johnson et al. (1975) test, if the ARM is more stable to AF demagnetization than

68

4 Laboratory Techniques

2

Detritai level c o n t a i n i n g M D m a g n e t i t e 9r i , i , i , i ' , ' i ' I' I ' I' I' I'

I

<E

o tt.

,- I

,

,,

1

2F"l -0.6

! , I , I , I , I , I Iil, 0.2 -0.4 -0.2 0

I,

lJl,

0.4

0.6

Magnetizing Field (T) Background sediment containing PSD magnetite

1

E

< ,~

0

-0.6

-0.4

-0.2

0

0.2

0.4

0.6

Magnetizing Field (T) Figure 4 . | | Hysteresis loops for U p p e r Pleistocene detrital layers from the L a b r a d o r Sea containing M D magnetite and intervening background sediment containing PSD magnetite (after Stoner et al., 1995a).

SIRM, the magnetite is considered to be predominantly in the SD grain size range (Fig. 4.13). Cisowski (1981) has shown that strongly interacting SD grains respond to the test, but that noninteracting or weakly interacting SD grains do not give the predicted response. Cisowski (1981) proposed a test for the degree of interaction based on the point of intersection (on the

4.4 Practical Guide to the Identification of Magnetic Minerals

69

0.6 -SD 0.5 PSD 0.4

--

0.3

-

0.2

-

0.1

-

1,1

9

0

1

1

2

3

--,-',w,,',,,i,

4

5

~'

6

OMD

7

Hrc/Hc Figure 4.12 Hysteresis ratios plotted on a Day et al. (1977) diagram, illustrating a mixing line in sediments containing PSD and MD magnetite (after Stoner et al., 1995a).

normalized moment axis) of the IRM acquisition and IRM demagnetization curves (Fig. 4.13).

f. Low-Temperature Treatment of IRM Transitions in the magnetic properties of magnetite, pyrrhotite, and hematite occur at low temperatures and they provide a potential means of magnetic mineral identification. Magnetite exhibits a crystallographic phase transition from cubic to monoclinic at 110-120 K. Associated with this transition, the anisotropy constant goes through zero as the easy axis of magnetization changes from [100] to [111] (Nagata et al., 1964). The so-called Verwey transition in magnetite can be recorded by a decrease in intensity of an IRM imposed at low temperature as it is allowed to warm through the transition temperature or alternatively as an IRM imposed at room temperature is allowed to cool through this temperature (Fig. 4.14). The transition has been clearly observed in the NRM iron ore samples (Fuller and Kobayashi, 1967) and the SIRM of European Mesozoic limestones (e.g., Heller, 1978; Mauritsch and Turner, 1975; Channell and McCabe, 1994) and deep-sea sediments (e.g., Stoner et aL, 1996, Fig. 4.14). The transition should be manifest as a relatively sharp change in magnetization intensity and should not be confused with the steady decrease in low-temperature IRM intensity due to unblocking of superparamagnetic grains. It was considered that the Verwey transition was restricted to MD grains (Fuller and Kobayashi, 1967); how-

70

4 Laboratory Techniques

1~

--

IRM acquisition

I

=

IRM demagnetization

'

= ' 20

0.8 r-

0.6 "o N (II

E

0.4

Z

0.Z

0

0

"

'

'

'

I 40

.

.

. 60

.

80

100

Applied Field ( m T ) Figure 4 . 1 3 The point of intersection of IRM acquisition and IRM demagnetization for a sample of white pelagic limestone is close to 0.5, consistent with noninteracting SD magnetite. According to Cisowski (1981), noninteracting or weakly interacting SD grains do not give the predicted response to the Johnson et al. (1975) test based on the relative AF stability of ARM and IRM.

ever, it is now clear that the transition occurs in near-stoichiometric magnetite crystals for all sizes above 37 nm, below which the transition is masked by superparamagnetic unblocking (Ozdemir et al., 1993). Small deviations from stoichiometry in magnetite, including oxidation to maghemite, have the effect of suppressing the Verwey transition (Ozdemir et al., 1993). The Morin transition in hematite at -10~ due to a change in crystallographic orientation of the spin axis, is more elusive. It has been observed in iron ore samples (Fuller and Kobayashi, 1967) but not, as far as we know, in sediments. Apparent observation in lake sediments (Creer et al., 1972) has been reinterpreted as due to disruption of grain orientation as a result of ice formation in unconsolidated sediment (Stober and Thompson, 1979). The pyrrhotite transition at 34 K has been recognized in igneous and metamorphic rocks (e.g., Dekkers et al., 1989).

g. Continuous Thermal Demagnetization of IRM Continuous measurement of remanence during heating has several advantages over stepwise thermal demagnetization. There is less likelihood of mineralogical changes, and no PTRM or VRM will be acquired by the

4.4 Practical Guide to the Identification of Magnetic Minerals

7]

1 9

4

0.8 nu

"O (1)

0.6

N

,.m

E tm

0.4

hemi~eiagic

o z

. . . . .

- -

0.2

- BG

:

Layers _ _ . . . . . . .

........

129 cm

BG 6 1 9 cm

~

'

'

'

1

20

'

'

'

I

'

'

40

1

60

'

..... - ~.... ~, . + ....~_ .+. _ ~

\ ' ~ " ' 4 - . - -1-- ,. ,++. .

LDC2

~

.

.

J

DC 4 LDC5

:

0

~..'~..0

'

'

I

'

80

Temperature

'

'

I

'

100

'

'

1

120

'

'

'

I

140

'

'

'

160

(~

Figure 4.14

Verwey transitions in magnetite-bearing Upper Pleistocene sediments from the Labrador Sea. Coarser grained detrital layers with higher proportion of MD magnetite show larger remanence drop than PSD-dominated background sediments. IRM acquired in 500 mT field at 10 K (after Stoner et aL, 1995a).

samples. Cryogenic magnetometers have not been used on a routine basis for continuous thermal demagnetization due to problems of construction of a heating system that does not interfere with the S Q U I D sensors. The most commonly used magnetometer for continuous thermal demagnetization is a modification of the Digico magnetometer (e.g., Heiniger and Heller, 1976). The loss of sensitivity of the larger fluxgate ring, which provides room for the furnace, relative to the room temperature ring (---factor of 10) is such that, for most sediments, the demagnetization experiments must be carried out on the stronger IRM rather than on the NRM. Continuous demagnetization of superimposed IRM components can be useful for determining the maximum blocking temperatures of two IRM components. h.

Susceptibility/Temperature

Initial susceptibility is often routinely measured after each measurement step during thermal demagnetization in order to monitor changes in magnetic mineralogy during heating. The maghemite-to-hematite transition at about 300~ produces a decrease in susceptibility which is a useful means of identifying the presence of this mineral. The susceptibility of maghemite is 2 to 3 orders of magnitude greater than that of hematite. This susceptibility

7~

4 LaboratoryTechniques

decrease will be masked if iron sulfides or titanohematite are present, because heating in air oxidizes the iron sulfides and titanohematites to magnetite, which has high susceptibility, comparable to that of maghemite. The formation of magnetite during the thermal demagnetization procedure can result in the development of spurious magnetizations. Note that the most common iron sulfide, pyrite, is paramagnetic and cannot therefore carry remanence.

i. Coercivity/Temperature An additional method for monitoring the change in mineralogy during heating involves the change of remanence coercivity (Her). The remanence coercivity of maghemite (and magnetite) is 1 to 2 orders of magnitude less than that of hematite. The maghemite-to-hematite transition at about 300~ will be accompanied by a change in remanence coercivity. The measurements of ncr can be carried out after each heating step with an alternating force gradient magnetometer (/zMAG) but the measurement procedure is considerably slower with more standard paleomagnetic equipment. For example, after each heating step, the sample is given 1 T IRM, and a back field IRM (antiparallel to the 1 T IRM) is then imposed in steps of 0.1 T. The back field corresponding to zero IRM corresponds to the remanence coercivity (Hcr). The increase in remanence coercivity in the 300-400~ temperature range has been cited as evidence for the presence of maghemite in red pelagic limestones (Channell et aL, 1982a).

j. Separation Techniques Magnetic mineral separation techniques are invariably selective, and not fully representative of the grain size and composition of magnetic minerals present in the sample; however, magnetic separation may be necessary for SEM/TEM, X-ray, Mossbauer, or chemical analyses. The importance of fine SD grains as remanence carriers emphasizes the necessity of making the separation technique as sensitive as possible to the fine grains. As grain shape has become an important criterion for distinguishing detrital and biogenic magnetite, separation procedures to prepare representative extracts for SEM and TEM observation have become more important. A recommended procedure which has been successful for extracting magnetite (including the SD fraction) is as follows: (i) Crush the sample (if necessary) in a jaw crusher with ceramic jaws. (ii) Use a mortar and pestle to produce a powder. (iii) Dissolve carbonate with 1N acetic acid buffered with sodium acetate to a pH of 5, changing the reagent every day until reaction ceases (several weeks). (iv) Rinse the residue with distilled water. (v) Agitate the

4.4 Practical Guide to the Identification of Magnetic Minerals

73

residue ultrasonically in a 4% solution of sodium hexametaphosphate to disperse the clays. (vi) Extract the magnetic fraction using a high-gradient magnetic separation technique (Schulze and Dixon, 1979), or alternatively, pass the solution (several times, if necessary) slowly past a small rare earth magnet.

Fundamentals of Magnetic Stratigraphy

5.1 Principles and Definitions The purpose of magnetic stratigraphy is to organize rock strata into identifiable units based on stratigraphic intervals with similar magnetic characteristics. Short-period behavior of the magnetic field (secular variation) can be useful for correlation, particularly in Holocene lacustrine sediments and archaeomagnetic studies (Ch. 13). At the other end of the geologic time scale, paleomagnetic pole positions have been used as a means of correlating Precambrian rocks (Irving and Park, 1972). Relative geomagnetic paleointensities as well as a wide range of nondirectional magnetic properties can be used for correlation (Ch. 14). The magnetic characteristic most often used, however, is the polarity of magnetic remanence. The polarity is said to be "normal" (north-seeking magnetization gives a northern hemisphere pole, as today) or "reverse" (north-seeking magnetization gives a southern hemisphere pole). When correlated to the geomagnetic polarity time scale (GPTS), the polarity record observed in stratigraphic sequences allows them to be placed in a chronostratigraphic framework. The most important geomagnetic property for stratigraphic purposes is aperiodic polarity reversal of the geomagnetic dipole field. The direction of magnetization of a rock is by definition its north-seeking magnetization, and as explained above, the magnetization can usually be designated either normal or reverse polarity. A problem may arise in Early Paleozoic or older rocks, or in displaced terrains, when it is unclear in which hemisphere the rocks acquired their magnetization. An ambiguity may be introduced as to the true polarity of the rock. In this case, an apparent polar wander path (APWP) must be established for the individual tectonic unit being

5.2 Polarity Zone and Polarity Chron Nomenclature

75

investigated. For instance, the North American APWP crosses the equator in Lower Paleozoic time, introducing some ambiguities as to which is the north pole and which is the south pole, particularly in regions where there has been large-scale local tectonic rotation. In practice, the polarity can be determined by following the motion of the pole sequentially back in time from the present. The terminology of magnetic polarity stratigraphy has been codified by the lUGS subcommission on the Magnetic Polarity Timescale (Anonymous, 1979). This body introduced the following useful definitions. 1. Magnetic stratigraphy or magnetostratigraphy~the element of stratigraphy that deals with the magnetic characteristics of rock units. 2. Magnetostratigraphic classification--the organization of rock strata into units based on variations in magnetic character. 3. Magnetostratigraphic units (magnetozones)~bodies of rock strata unified by similar magnetic characteristics which allow them to be differentiated from adjacent strata. 4. Magnetostratigraphic polarity classification~the organization of rock strata into units based on changes in the orientation of remanent magnetism in rock strata, related to changes in the polarity of the Earth's magnetic field. 5. Magnetostratigraphic polarity reversal horizons are surfaces or very thin transition intervals in the succession of rock strata, marked by changes in magnetic polarity. In practice, the term magnetostratigraphic polarity reversal horizon may be allowed to stand for a finite transition interval. Less commonly, the polarity change takes place through a substantial interval of strata, and the term magnetostratigraphic polarity transition zone should be used. Magnetostratigraphic polarity reversal horizons and magnetostratigraphic polarity transition zones may be referred to simply as polarity reversal horizons and polarity transition zones, respectively, if in the context it is clear that the reference is to changes in magnetic polarity. Polarity reversal horizons or polarity transition zones provide the boundaries for polarity stratigraphic units. 6. Magnetostratigraphic polarity units (or zones) are bodies of rock strata, in original sequence, characterized by their magnetic polarity, which allows them to be differentiated from adjacent strata.

5.2 Polarity Zone and Polarity Chron Nomenclature The basic magnetostratigraphic unit is the magnetostratigraphic polarity zone (in accordance with the nomenclature recommended in the Interna-

76

5 Fundamentals of Magnetic Stratigraphy

tional Stratigraphic Guide, pp. 8-14, see Anonymous, 1979). A magnetostratigraphic polarity zone may be referred to simply as a magnetozone, if in the context it is clear that the reference is to magnetic polarity. Polarity zones are bounded above and below by polarity reversal horizons or polarity transition zones (see definition 5, above). Polarity zones may be of different rank. Magnetostratigraphic polarity zones may (1) consist of strata with a single polarity throughout or (2) be composed of an intricate alternation of normal and reverse units or (3) be dominantly either normal or reverse magnetozones, but with minor subdivisions of the opposite polarity. Thus a zone of dominantly normal polarity may include lower rank units of reverse polarity. The words normal and reverse, where it is possible to indicate the polarity unambiguously, may be included as a part of the term name as an aid to clarifying its significance. However, as mentioned above, in premiddle Paleozoic rock sequences, it is sometimes difficult to determine what is normal and what is reverse. Also, in rare circumstances, excursions of the paleofield may provide polarity units that cannot be classified as normal or reverse, but may be designated as intermediate in direction. The end of a successful magnetostratigraphic survey is the identificat i o n o f a series of normally and reversely magnetized polarity zones of varying length arranged in stratigraphic order. At this stage, it is useful to give each magnetozone (polarity zone) a number/letter and geographic designation, such as, from base of the section: San Pedro N1, R1, N2, R2, etc., or Gubbio A - , B+, C - , D + , etc. At Gubbio (Alvarez et al., 1977; Napoleone et al., 1983), polarity zones were designated in ascending alphabetical order, normal zones distinguished by " § and reverse zones by " - " . The thinner polarity zones are given lower hierarchy and are designated as subzones. At Gubbio, subzones within D + are labeled D I + , D 2 - , D3+. The labeling of polarity zones (magnetozones) in individual sections (Fig. 5.1) is useful as it is agnostic and separates the observations from the correlation to the geomagnetic polarity time scale. It is a mistake to confuse polarity zone labels with oceanic magnetic anomaly numbers; for example, a polarity zone should not be referred to as "anomaly 34." Oceanic magnetic anomalies and polarity zones are distinct phenomena, and they should not be confused. They may, however, have a common cause and be time correlative.

77'

5.2 Polarity Zone and Polarity Chron Nomenclature

MAGNETIC POLARITY ZONES

VG P Latitude (o) - 90

0

t.o h_

,90

Gubbio H+ GF3+ F2FI+

E Tertiary 349. 347.6 343 Cretaceous 336.2 335.9 325

E.

300 -

D3+ D2DI+ C-

305 302 299 289 282

E

v

(D

O C

B+

03

0~ s o c-{3. 03 133 ,

m

.

m

.

m

03

i,,_ .,-, 03

219 A200 -

197

tO

N

E

normal

O

z

c" O .J

reverse

O .Q

100 l:igure 5.l Magnetic stratigraphy of the Upper Cretaceous Scag]ia Rossa Fro. at Gubbio (italy). The VGP latitude is presented for each sample. The polarity zones are labeled alphabetically (after Alvarez et aL, 1977).

78

5 Fundamentals of Magnetic Stratigraphy

The recommended hierarchy in magnetostratigraphic units and polarity chron (time) units is as follows:

Magnetostratigraphic polarity units Polarity Polarity Polarity Polarity Polarity Polarity

megazone superzone zone subzone microzone cryptochron

Geochronologic (time) equivalent

Chronostratigraphic equivalent

Approximate duration (yr)

Megachron Superchron Chron Subchron Microchron Cryptochron

Megachronozone Superchronozone Chronozone Subchronozone Microchronozone Cryptochronozone

108-109 107-108 106-107 105-106 <105 uncertain existance

Unlike most branches of stratigraphy, the chronologic system in use in magnetic stratigraphy was determined simultaneously with its development. Initially, Cox et al. (1963) boldly divided the reversal sequence from 0 to 4 Ma into four magnetic epochs (chrons): Brunhes, Matuyama, Gauss, and Gilbert, named after great geomagnetists. It is noteworthy that, due to techniques used in determining polarity and dating individual lava flows, type sections for the Brunhes, Matuyama, Gauss, and Gilbert were never established. Type localities are known, however, for subchrons such as the Jaramillo, at Jaramillo Creek in New Mexico. In present usage, polarity chron labels, Brunhes, Matuyama, Gauss, and Gilbert, are used for the Pliocene and Pleistocene, with the remainder of the Cenozoic being subdivided into polarity chrons designated by numbers correlated to oceanic magnetic anomalies. For example, polarity chron C26 encompassed the time represented by magnetic anomaly 26 and the reversed interval preceeding it (Fig. 5.2). The numbered magnetic anomalies in the Late Cretaceous and Cenozoic correspond to intervals of normal polarity of the geomagnetic field. The polarity chron nomenclature has evolved progressively to accommodate additional polarity chrons (LaBrecque et al., 1977; Harland et al., 1982; Cande and Kent, 1992a). For example, the three polarity intervals within C23n, from old to young, are labeled C23n.2n, C23n.lr, and C23n.ln (Fig. 5.2). The Cande and Kent (1992a) chron nomenclature (Fig. 5.2) is slightly modified and expanded from that of LaBrecque et al. (1977). As an example, let us examine chron 3A (Fig. 5.3) which is divided into a dominantly normal polarity upper segment

Figure 5.2 Magnetic polarity time scale and polarity chron nomenclature for Campanian to Eocene time (after Cande and Kent, 1992a).

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5 Fundamentals of Magnetic Stratigraphy

al r

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Figure 5.3 Magnetic polarity time scale and polarity chron nomenclature for Late Miocene to recent (after Cande and Kent, 1992a).

(3An) and a preceeding reverse interval (3Ar). The upper normal/reverse interval is designated as 3An.1. The reverse interval would then be 3An.lr. In this way a complete designation of the time scale can be achieved. The location of a geological occurrence in time within such a system may be precisely placed using the chron label with suffix "x", where

5.2 Polarity Zone and Polarity Chron Nomenclature

x = (Xe

and

-

Xco)/Xcy-

13|

Xco)

Xe = stratigraphic position of event S c y -- location of younger chron boundary Xco - location of older chron boundary,

such that 3An.1r(0.25), for example, is a location within the subchron, 25% from the base of the subchron. For the Late Jurassic and Early Cretaceous, Harland et al. (1982) labeled polarity chrons according to the correlative M-sequence oceanic magnetic anomaly. In contrast to Cenozoic marine magnetic anomalies, most but not all of the Mesozoic (M-sequence) oceanic magnetic anomalies correlate to reverse polarity chrons. Intervening normal polarity chrons are given the number of the next (older) reverse chron with "n" appended. As for the Cenozoic, we recommend use of the prefix "C" when the label refers to a polarity chron (time interval) rather than an oceanic magnetic anomaly. The most complete record of the reversal pattern of the geomagnetic field since 160 Ma is preserved in the crust of the ocean and detected by seaborne/airborne magnetometer surveys. These magnetic anomalies are interpreted to be due to reversals of the geomagnetic field recorded in the rocks of the ocean floor. The sea floor in a major ocean is quasistratigraphic in that it gets younger toward the spreading axis; however, it is not in superposition in the usual sense. The reversal record as preserved in the sea floor has become the template (a sort of type section) for geomagnetic reversal history of the last 160 My and is the basis for the geomagnetic polarity time scale (GPTS). The designation of type sections, in the classical stratigraphic sense, for the geomagnetic reversal pattern since 160 Ma is unnecessary. On the other hand, type sections for magnetobiochronology, the correlation of the biologic record to the magnetic polarity sequence, might be desirable. For example, a potential type section for the Paleogene might be the Contessa section in Italy (Lowrie et aL, 1982). The type section concept for polarity history is the only way in which the polarity record can be built up for time intervals not recorded by present-day sea floor. For periods prior to 160 Ma, for which no oceanic magnetic anomalies are available, the establishment of polarity sequences must be carried out using classical stratigraphic principles involving type sections. An example of a viable magnetostratigraphic type section is that recording the Kiaman Superchron (Interval) in Australia, where the superchron is thought to be tied directly to the rock record (Irving and Parry, 1963).

82

5 Fundamentals of Magnetic Stratigraphy

5.3 Field Tests for Timing of Remanence Acquisition Undoubtedly, one of the most important technical challenges in any paleomagnetic research is to determine the age of magnetization components. It is of utmost importance to determine which, if any, of the magnetization components was acquired near the time of deposition of the sediments or cooling of the igneous rock. In marine sediments, experiments have shown that the detrital remanence (DRM) is acquired at some position below the sediment-water interface, where the water content falls to a level at which the magnetic grains can no longer rotate (Irving and Major, 1964; Kent, 1973; Verosub, 1977). For pelagic sediments, a reasonable estimate of the depth at which the DRM becomes locked is 10-20 cm below the sediment-water interface (3.3c). The age of the sediment magnetization will not, therefore, coincide with the age of sediment deposition but will lag behind by intervals dictated by the lock-in depth. In marine sediments retrieved from the bottom of the ocean, the stability of the record can be established by laboratory demagnetization experiments, by the internal coherence of the reversal pattern, and by a reversal test. A reversal test is based on the fact that normal and reverse intervals of the core should yield mean directions 180~ apart (Fig. 5.4). This is an important test since any magnetization component acquired later will tend to affect the reverse and normal directions differently. It is usual to test whether the circles of confidence at the 95% level overlap for the reverse and normal mean directions. The results of this test depend on sample size and dispersion of the data, and if one of the sets of directions has a large

N

9Mean of reverse sites flipped through 180 ~ x Mean of normal sites § Dipole field at sampling site Figure 5.4 Mean directions of normally and reversely magnetized sites from Curtis Ranch, San Pedro Valley (Arizona) with overlapping circles of 95% confidence. Reverse mean direction is flipped through 180~ (after Johnson et al., 1975).

5.3 Field Tests for Timing of Remanence Acquisition

~

O995 then no significant difference will be found. McFadden and McElhinny

(1990) have proposed a simple version of the reversal test in which the mean directions of the reversed and normal data sets and a95 are utilized. It is assumed that both sets have a common precision. In the case where both sets of data have N greater than 5, the following equation can be used to calculate the critical angle Tc with p = 0.05, where N = N1 + N2 and R1 and R2 are the vector sums for the two polarities. 1

If the angle To between the two mean directions is greater than the angle Tc then the hypothesis of a common mean direction may be rejected at the 95 percent confidence limit and the reversal test fails. The "positive" reversal tests are classified on the basis of Tc, as A if Tc -< 5~ as B if 5 ~ < Tc <10 ~ as C if 10 ~ -< Tc -< 20~ Very few data are classified as A since the requirements are so stringent. The data presented in Fig. 5.4 are classified as B, since the angle T~ is greater than 10 ~ and less than 20 ~ If Tc exceeds 20 ~ the reversal test is considered to be negative. Sedimentary rocks often become involved in tectonic processes and are then exposed through erosion at the Earth's surface. This makes possible other tests for stability such as the bedding tilt test suggested by Graham (1949). Graham pointed out that if the direction of remanent magnetization was stable from the time of rock formation, the direction of remanence determined from tilted rock strata would be dispersed but would coincide after correction for bedding tilt (Fig. 5.5). It is now well known that directions of magnetization can be acquired before, during, and after the rocks have been deformed. It is therefore necessary to determine the stage in the tilting process at which a particular magnetization component reaches maximum concentration. This is usually done by incrementally untilting the beds and calculating the Fisher (1953) concentration parameter (k) at each step. The stage in this process at which the Fisher "precision" or concentration parameter (k) is maximized is usually taken to represent the attitude of the beds at the time that the magnetization component was acquired. In Fig. 5.5, it can be seen that the highest values of k for component A occur when the beds are in their present position (0% untilting). Therefore, component A was acquired after folding of the rocks. The B component of magnetization reaches its greatest concentration when the beds are unfolded by ---50%; this component was therefore acquired during the deformation process and is a synfolding magnetization. Component C reaches its highest concentration after 100% unfolding. The C component was therefore acquired before folding of the beds and is the only component which can be used for magnetostratigraphic

84

i.,i .i, i: ."

" "i- i. i..".~i..:.

5 Fundamentals of Magnetic Stratigraphy

.i

~;~

i"

c

% unfolding

Figure 5.5 Schematic presentation of fold and conglomerate test. Three magnetization components are present in these sediments. The A component is postfolding and was acquired with the sediments in their present position. Component B is synfolding and was acquired when the rocks were tilted by 50% of their present attitude. Component C is prefolding and the directions group best after restoring the beds to horizontal. The magnetization directions in syndepositional rip-up clasts show randomly directed C components of magnetization, indicating that the C component was acquired very early in the history of the sediment, k is the Fisher (1953) concentration parameter.

purposes. The improvement in grouping may be tested statistically (McElhinny, 1964). In this test, the precision of the mean direction before tilt correction (kl) is compared to the precision after tilt correction (k2). The ratio k2/kl can be compared with difference ratio tables with equal degrees of freedom (2N - 1). This test has been widely applied by paleomagnetists for many years and is still in use. McFadden and Jones (1981) pointed out that the McElhinny (1964) tilt test was invalid because the two estimates of k are related through the bedding attitude. The McElhinny (1964) fold test is often said to be conservative; however, it is clear that in some cases it can be misleading, particularly when the magnetization being tested was acquired during folding (e.g., Kent and Opdyke, 1985). McFadden (1990) proposed a fold test based on correlations between the distribution of magnetization directions and bedding attitudes. Recently two new tilt tests have been proposed that have considerable merit. Watson and Enkin (1993) argued that the problem could be solved by parameter estimation. This method legitimizes the use of Fisher's precision parameter (k). Their method involves parametric resampling of the original

5.4 Field Sampling for Magnetic Polarity Stratigraphy

85

data and plotting k as a function of the degree of untilting (Fig. 5.6). The number of samples, observed mean direction, and k are used to create synthetic data sets which are sequentially untilted around a horizontal axis generating, say, 1000 synthetic curves. The technique permits the calculation of the best estimate of the amount of unfolding at the time of magnetization and a calculation of 95% confidence limits for this value. Another tilt test has been proposed by Tauxe and Watson (1994) which is based on eigenvector analysis and parameter estimation techniques. This fold test has the advantage that the polarity of the resultant vector is not critical to the analysis, so no assumptions are required as to polarity or whether the data have a Fisherian distribution. A bootstrap method is employed to determine 95% confidence limits on parameters. The conglomerate test utilizes the fact that an early magnetization in conglomerate clasts will be randomly distributed (Fig. 5.5). The most powerful conglomerate test is that applied to syndepositional conglomerates, such as rip-up clasts. These synsedimentary conglomerates can be formed by river channels meandering into floodplains containing previously deposited muds and silts. In some cases coherent blocks of sediment are incorporated into the sandstone of an ancient river channel. If magnetization directions in these clasts have magnetic properties similar to those of the parent sediment and possess a stable component which is randomly directed, then the investigator has demonstrated that the component in question was acquired early in the history of the rock (Fig. 5.7). The randomness of the clast magnetization directions can be tested using the Watson (1956b) test for randomness. Another important test of early acquisition of magnetization (Graham, 1949) is the "slump" test. Syndeposition slumps are often present in rock sequences; however, application of this test is uncommon. If magnetization was acquired during deposition, remanence directions from slumps should be different from the unslumped parent rock. Alvarez and Lowrie (1984) have demonstrated that Upper Cretaceous pelagic limestones in Italy contain large slumps that yield directions which deviate widely from those in the unslumped parent rock, indicating that these sediments were magnetized early in their history.

5.4 Field Sampling for Magnetic Polarity Stratigraphy The geomagnetic field reverses in an aperiodic fashion, with intervals of constant polarity having durations ranging from 20 ky to 50 My. An ideal sampling scheme would be one that is designed to monitor the geomagnetic field every ---5000 years. This is often impractical or impossible due to the

a

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0

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0

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50

100

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Figure 5.6 The value of the Fisher (1953) concentration parameter (k) as a function of the

degree of untilting for 20 examples of parametric resampling of the original data. The histogram shows the distribution of the maxima from 1000 trials in groups of 5% ranges of unfolding. (a) Data from Gough and Opdyke (1963). (b) Data from Stamatakos and Kodama (1991).

5.4 Field Sampling for Magnetic Polarity Stratigraphy

N

87

N

W

E

S

S

Figure 5.7

Positive conglomerate test on Carboniferous Mauch Chunk sediments from Pennsylvania. Magnetization directions from individual clasts indicate a C component predating deposition of the conglomerate and a B component postdating the deposition of the conglomerate.

type of rock being sampled. In sampling marine sedimentary cores, it is common practice to remove samples at 10-cm intervals by sawing cubes if the sediments are dry, or by pressing plastic containers into the sediment if the samples are wet. Sampling at this density for constant pelagic sedimentation rates of 10 m/My would resolve polarity chrons with duration greater than ---10 ky. Usually, sample volume is about 7-10 cm 3. The sensitivity of modern magnetometers is such that much smaller samples could be measured; however, small samples are more difficult to orient and handle. In the study of polarity transitions, thin 3-mm slices of sediment have been used (e.g., Clement et al., 1982). The time between reversals of the geomagnetic field is variable. Polarity subchrons as short as 20 ky in duration are present in the polarity record (e.g., Cobb Mountain and R6union subchrons). Ideally, any sampling scheme should attempt to sample intervals of this duration. At pelagic sedimentation rates, such resolution would require continuous back-to-back sampling with 7-cm 3 plastic cubes. Pass-through cryogenic magnetometers can be used to measure split half-cores or whole cores, without subsampling. One drawback of continuous measurement is that it is often desirable to do progressive demagnetization using higher alternating fields than are available in these instruments, or to do thermal demagnetization. In these cases, discrete and separate samples are required. A pass-through magnetometer, which can produce independent measurements at ---20 cm spacing, has been used for routine core analysis aboard the ODP vessel R/V Joides Resolution since

~8

5 Fundamentalsof Magnetic Stratigraphy

1983. Weeks et al. (1993) describe a modified 2G-755R magnetometer in which the SQUID sensing coils are arranged for high resolution measurement of magnetization from continuous "u-channels" tracked through the sensing region. The width of the response function of the pick-up coils is about 3.5 cm. Deconvolution of the signal in the time domain (Constable and Parker, 1991) can give comparable resolution to that achieved by back-toback discrete sampling. The u-channel (see Tauxe et al., 1983b; Nagy and Valet, 1993), typically 1.5 m in length with a cross sectional area of about 4 cm 2, allows continuous sampling of the core with less sediment distortion than for discrete back-to-back sampling. The rate of sediment accumulation is generally variable through time. Sadler (1981) emphasized this variability and termed the process inherently "unsteady," characterized by large changes in rates of accumulation over short time periods. In some environments, measured rates of sediment accumulation during floods may be ---102 cm/hour, whereas accumulation rates integrated o v e r 10 6 years would give much lower values, implying large changes in rates of sedimentation, abundant hiatuses, and unconformities. The ideal condition for a continuous record of geomagnetic polarity would be a continuous rain of sedimentary particles accumulating uninterrupted through time. Such ideal conditions are rare, but in certain cases apparently continuous sediment accumulation has been recorded for marine sediments over significant periods of time (5-10 My) and for lacustrine sediments over intervals of ~ 1 0 4 years. Variable rates of sedimentation caused by tectonics, or changes in climate or paleoceanographic circulation, are well known. When conditions are favorable, complete and continuous reversal sequences have been recorded in marine cores (Opdyke et aL, 1966; Foster and Opdyke, 1970; Tauxe et al., 1983c) and in outcrop (Lowrie and Alvarez, 1977b; Channell and Erba, 1992). Some marine sediments preserve such excellent records of the geomagnetic field that even shortperiod behavior of the field, such as the process of field reversal, can be studied (Opdyke et al., 1973; Clement and Kent, 1984). Lacustrine sediments can also be excellent recorders of the geomagnetic field and they have often been used to study secular variation and magnetic excursions (Ch. 13). Unlike deep-sea sediments, terrestrial sediments are rarely, if ever, homogeneous with respect to lithology or sedimentation rates. Terrestrial sections are usually selected for sampling for particular reasons, such as the presence of important vertebrate faunas, radiometrically dated horizons, or climatically indicative levels in loess, tills, or pollen-rich deposits. In the Siwaliks of Pakistan and India, the sedimentary rocks comprise mudstones, sandstones, and gravels deposited in the subsiding foreland basin in the front of the rising Himalayan mountain chain (Opdyke et al., 1979; Johnson et al., 1982, 1985; Tauxe, 1980). The sediments were deposited by meandering river systems across the floodplain and consist of fining upward se-

5.4 Field Sampling for Magnetic Polarity Stratigraphy

89

quences of river channel sands and overbank silts and muds on which soils developed. Such a sequence must contain many disconformities. An analysis of the variability of sedimentation in this facies indicates that time gaps of thousands of years are present (Friend et al., 1989; McRae, 1990a,b). Nevertheless, Johnson et al. (1985) have shown that over periods of the order of 105 to 106 years the magnetic polarity record is amazingly complete. Sampling in the Siwaliks foredeep sediments was confined to the finegrained mudstones and siltstones, with sandstone intervals not sampled. The sites were, therefore, distributed at random intervals determined by lithology and outcrop availability, with stratigraphic spacing of sites of 5 to 10 m. Studies on sediments in Argentina deposited by unchannelized flows on sandflats and distal alluvial fans in an intermountain basin indicate that the record is fairly complete on time scales greater than 104 years, although in this case the stratigraphic spacing of sites was --25 m (Beer, 1990). Most investigations of intermontane fluvial systems have shown them to be faithful recorders of the geomagnetic field for features with duration greater than ---105 years, although obviously with lower rates of sedimentation shorter sampling intervals are required. Several studies in Pakistan and Argentina have dealt with the variability and reproducibility of polarity zones along strike (Behrensmeyer and Tauxe, 1982; Johnson et al., 1988; Beer, 1990). Johnson et al. (1988) have shown that individual polarity zones (Fig. 5.8) may vary along strike by up to a factor of two in this type of sedimentary environment. It is a great advantage in magnetostratigraphic studies if sections of appreciable thickness (102-103 m) are available for sampling. Thick sections increase the probability that more than one polarity zone is present in the section and increases the chances that the reversal pattern can be correlated to the GPTS. The following technique for sampling unindurated sediments was developed by Johnson et al. (1975) for study of Plio-Pleistocene nonmarine sediments of the San Pedro Valley of Arizona. During the initial field work a suitable sampling interval is chosen. Sites are preferentially selected in finer grained sediments (silt size or finer) at stratigraphic spacing of 2-20 m. Sections are often sited in arroyos where erosion has exposed fresh unweathered material well away from local peaks or promontories which might be the sites of lightning strikes. At individual sites, it is often necessary to excavate a pit to obtain unweathered material, using a pickmattock, spade, or some other suitable implement. After unweathered material is exposed, a horizontal or vertical face is fashioned on a fist-sized sample, using a hand rasp. The strike and dip of the surface are measured and marked with fine pen or pencil. Three or more independently oriented samples are generally chosen at each site, so that the polarity can be determined unambiguously and rudimentary site statistics can be deter-

90

5 Fundamentals of Magnetic Stratigraphy

o

2

~

DISTANCE ALONG STRIKE (km)

?

8

lj0

12

14

Isochron= 12.74 M a

+80-

g. +40E

.J uJ > u,I ._1

0-

,1,a. '~ n- -40O

-80-

-120K3 HN

K2

CP

K4

Figure 5.8 Variation of the thickness of polarity zones along 15 km of strike in Miocene molasse sediments in Pakistan. An individual polarity zone may vary in thickness by a factor of 2. Time scale is from Mankinen and Dalrymple (1979) (after Johnson et al., 1988).

mined. After the sample is returned to the laboratory it is fashioned into a cube (approximately 2.5 cm on a side) using a band saw or grinding wheel. Although the technique described above is laborious, once the cubes have been fashioned they are easy to handle and measure. Unindurated sediments are often sampled by pressing the samples into plastic boxes or cylinders in the field. The technique is rapid and is convenient for sample transportation and handling. If thermal demagnetization is necessary, the samples must be dried and/or indurated, removed from the plastic containers, and wrapped in aluminum foil for heating and measurement. Lava flows or indurated sediments in outcrop can be drilled using a portable gasoline-powered drill. Pelagic limestones in land section, such as those at Gubbio (Italy), are usually sampled by drilling individual oriented cores (2.5 cm diameter, 5-10 cm in length) spaced at 20-50 cm stratigraphic intervals. In lava sequences, three or more cores are usually taken from each lava flow. The most intensive study of this type has been performed by Watkins and Walker (1977) on the Icelandic lavas.

5.6 Correlation of Polarity Zones to the GPTS

91

5.5 Presentation of Magnetostratigraphic Data Magnetostratigraphic data can be presented as plots of declination and inclination of the primary magnetization components as a function of stratigraphic position in the section. Magnetization directions from individual stratigraphic horizons are often represented as virtual geomagnetic poles (VGPs) computed using the dipole formula (Ch. 2). VGPs from individual samples/horizons are compared with the overall paleomagnetic pole computed from the mean magnetization direction for the entire section. Variation in VGP in the section is then expressed as a VGP latitude which represents the latitude of individual VGP in the coordinates of the overall paleomagnetic pole. VGP latitudes close to + 90 ~ represent normal polarity, and VGP latitudes close to - 9 0 ~ represent reverse polarity (Fig. 5.1). VGP latitudes are plotted as a function of stratigraphic position of the samples/ horizons. The polarity interpretation is usually presented as black (normal) and white (reverse) bar diagrams.

5.6 Correlation of Polarity Zones to the GPTS The investigator is often confronted with a series of normal and reverse polarity zones, the pattern of which may correlate with more than one segment of the GPTS. Usually, it is initially assumed that the sedimentation rate, or rate of extrusion of the lavas, is more or less constant, and trial correlations are made on the basis of the fit of the observed polarity pattern to the GPTS. In practice, correlations to the GPTS are aided by radiometric ages or biostratigraphic events in the section, which have been correlated to the GPTS elsewhere. For example, in the Haritalyangar section in India (Fig. 5.9), the correlation of polarity zones to the GPTS is aided by magnetostratigraphic data from Pakistan where some of the same Miocene vertebrate fossils are correlated to radiometric ages (on volcanic ashes) and hence to the GPTS. The fossil assemblage at Haritalyangar is equivalent to that in Pakistan which is correlated to the chron 3A to chron 4A interval, and on this basis the reversal sequence can be unequivocally correlated to the GPTS on the basis of pattern fit (Fig. 5.9). The resulting regression is significant at the 99% confidence level, allowing sediments in India to be securely correlated to the GPTS and to similar sediments in Pakistan. It also indicates that the rate of sedimentation is reasonably constant. The Haritalyangar section is ideal in that it is long and continuously exposed. Short sections which contain few polarity zones are much more

92

5 Fundamentals of Magnetic Stratigraphy

1600-

1400-

r 1200 (1) E

N5

I

. ..,,.

h..

600-

03

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or r

"r" o

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~

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o rr

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Ra -r O N3 Z

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200

0

'

1;1

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.

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.

.

.

7

(Ma)

Figure 5.9 Magnetostratigraphy of the Haritalyangar section in India and correlation to the GPTS of Mankinen and Dalrymple (1979). The high linear regression value (r = 0.998) indicates that the correlation to the GPTS is feasible and that the sedimentation rate was more or less constant (after G. D. Johnson et aL, 1983).

difficult to correlate. In some cases, prominent subchrons may be missing because of hiatuses. Spurious apparent polarity subchrons can be an artifact of a hard secondary magnetization not (completely) removed by the demagnetization techniques. Cross-correlation techniques can be used to test the correlation of polarity zones to the GPTS (e.g., Lowrie and Alvarez, 1984; Langereis et al., 1984). The polarity zone pattern in stratigraphic section can be compared to the GPTS and the correlation coefficient computed for each successive match. The minimum value of the correlation coefficient which is significant at the 95% confidence level can be used to screen possible correlations. The optimal correlation can then be chosen based on stratigraphic age constraints. It is important to note that this technique is most useful in sections which record multiple polarity chrons, and where sedimentation

5.7 Quality Criteria for Magnetostratigraphic Data

93

rates are fairly constant such that the polarity zone pattern reflects the polarity chron pattern in the GPTS.

5.7 Quality Criteria for Magnetostratigraphic Data Van der Voo (1990) has recently suggested a reliability index for paleomagnetic studies. The index is based on seven reliability criteria, and is designed for paleomagnetic studies that contribute to construction of apparent polar wander paths. We propose a reliability index for magnetostratigraphic studies based on ten criteria. Some of the criteria are the same as those proposed by Van der Voo (1990); however, some are unique to magnetostratigraphy. The criteria and their justification are outlined below: 1. Stratigraphic age known to the level of the stage, and associated paleontology presented adequately. The placement of fossil occurrences and ranges with respect to the magnetic polarity determinations is often critical to interpretation. 2. Sampling localities placed in a measured stratigraphic section. Unless the section is accurately measured, correlation of the polarity zone record to the GPTS will be compromised. 3. Complete thermal or alternating field demagnetization performed and analysis of magnetization components carried out using orthogonal projections (Wilson, 1961; Zijderveld, 1967). All modern studies employ demagnetization, and vector analysis allows the identification of magnetization components present in the samples. Blanket demagnetization at an individual temperature or alternating field value is no longer considered adequate. 4. Directions determined from line-fitting least squares analysis (Kirschvink, 1980). This type of analysis has been widely applied in recent years. 5. Numerical data published completely. Black-and-white bar diagrams are not adequate. Data are best presented as VGP latitude/stratigraphic distance plots and/or as declination/stratigraphic distance and inclination/ stratigraphic distance plots. The statistical parameters (e. g., Fisher, 1953) should be fully documented. 6. Magnetic mineralogy determined. It is often useful to know the magnetic mineralogy, which may constrain the timing of remanence acquisition. 7. Field tests for the age of the magnetization are presented (tilt tests, conglomerate tests). If possible, these should always be attempted since they constrain the timing of remanence acquisition.

94

5 Fundamentals of Magnetic Stratigraphy

8. Reversals are antipodal. The field should reverse through 180~. If this reversal test fails, the investigator probably has not fully removed the secondary magnetization components. 9. Radiometric ages are available in the stratigraphic sections. Radiometric dates, especially 4~ or U-Pb ages from volcanic ashes or bentonites, are very useful for correlating polarity zones to the GPTS. 10. Multiple sections. Magnetic stratigraphy is more convincing if multiple sections yield similar or overlapping polarity zone patterns. Quite clearly, few studies will be able to meet all these criteria since, for example, radiometric dating is often not possible, and field tests are impossible in flat-lying beds or deep-sea cores. However, ratings of at least 5 (out of 10) should be achieved by modern magnetostratigraphic studies.

The Pliocene-Pleistocene Polarity Record

6.1 Early Development of the Plio-Pleistocene GPTS The Plio-Pleistocene geomagnetic polarity time scale (GPTS) developed from the early studies of Cox et al. (1963a) and McDougall and Tarling (1963a) which coupled K-Ar age determinations and magnetic remanence data in volcanic rocks from California and the Hawaiian Islands. By the end of the 1960s, most of the Plio-Pleistocene polarity chrons had been identified in radiometrically dated volcanic rocks (Fig. 1.2). As more KAr age determinations became available the chronogram technique for optimizing the age of polarity chron boundaries was introduced (Cox and Dalrymple, 1967). The chronogram is a plot of standard deviation or error function of available ages close to a polarity chron boundary against trial ages for the boundary (see Harland et al., 1990, pp. 106-109). The minimum error function corresponds to the best estimate of the age of the polarity chron boundary. Subchrons within the Plio-Pleistocene were initially identified by their associated radiometric ages. It is important to note that the early studies of Plio-Pleistocene geomagnetic polarity coupled K-Ar and magnetic data from widely spaced locations and, apart from the work in Iceland (e.g., Watkins et al., 1977), were not carried out in continuous stratigraphic sections. Plio-Pleistocene polarity subchrons were initially documented in the following publications: Jaramillo (Doell and Dalrymple, 1966), Cobb Mountain (Mankinen et al., 1978), Olduvai (Gromm6 and Hay, 1963), Kaena and R6union (McDougall and Chamalaun, 1966), Mammoth (Cox 95

96

6 The Pliocene-Pleistocene Polarity Record

et al., 1963), Cochiti (Dalrymple et al., 1967), and Nunivak (Hoare et al.,

1968). The two oldest subchrons in the Gilbert were initially identified by Hays and Opdyke (1967) from marine cores and by Pitman and Heirtzler (1966) from marine magnetic anomalies. Eventually they were identified from lava sequences in Iceland and named the Sidufjall and Thvera subchrons (Watkins et al., 1977).

6.2 Subchronswithin the Matuyama Chron Subsequent to the recognition of the existence of a normally magnetized lava at 1.8 Ma by Gromm6 and Hay (1963) from Olduvai Gorge, McDougall and Wensink (1966) reported a series of polarity zones (from base: R-NR-N) in a sequence of lavas from Iceland with ages of 1.60 _+ .05 Ma. Due to the significant age difference between these normal polarity zones and that observed at Olduvai, the normal intervals from Iceland were considered to be distinct from the Olduvai subchron and were called the Gilsa subchrons. Magnetic strat!graphy in deep-sea cores and oceanic magnetic anomaly profiles, however, appeared to indicate a single normal polarity subchron in the lower middle part of the Matuyama Chron, closely coincident with the Pliocene-Pleistocene boundary. The question arose, therefore, as to whether this interval of normal polarity should be called the Gilsa or Olduvai subchron. This led to a great deal of confusion and controversy (Watkins et al., 1975) which was finally resolved in favor of the name Olduvai. Recently, however, a short-lived normal subchron above the Olduvai has been documented by Clement and Kent (1987) and was called by them the Gilsa subchron (Fig. 6.1). Two short intervals of normal polarity below the Olduvai were identified in sea floor spreading anomalies and called anomalies "x" and "y" by Heirtzler et al. (1968). The youngest of these normal intervals probably coincides with a normally magnetized lava dated by McDougall and Chamalaun (1966) at 2 Ma and named the R6union Event by Gromm6 and Hay (1971). This subchron has a duration of approximately 20 ky (Clement and Kent, 1987; Khan et al., 1988). Mankinen et al. (1978) detected a short interval of normal polarity preceding the Jaramillo which they called the Cobb Mountain subchron. This subchron has been observed in high sedimentation rate cores from the Caribbean Sea (Kent and Spariosu, 1983) and the North Atlantic (Fig. 6.1; Clement and Kent, 1987). Another geomagnetic excursion has been reported between the Jaramillo subchron and the Brunhes/Matuyama boundary by Maenaka (1983) and Mankinen and Gromm6 (1982) and it

6.2 Subchrons within the Matuyama Chron

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6.3 Magnetic Stratigraphy in Plio-Pleistocene Sediments

99

is thought to have occurred between 0.84 and 0.85 Ma (Champion et al., 1988). A subchron of comparable age was observed in a loess sequence from China by Heller and Liu (1984). In the time scale presented here (Fig. 6.2), only those subchrons recorded in oceanic magnetic anomaly records and in magnetostratigraphic section are included.

6.3 Magnetic Stratigraphy in Plio-Pleistocene Sediments The intensive magnetostratigraphic study of marine sediments of PlioPleistocene age began in the early 1960s utilizing the piston core collections at Lamont Doherty Geological Observatory (LDGO) and cores from the Eltanin collection stored at Florida State University (Opdyke et al., 1966; Ninkovitch et aL, 1966; Opdyke and Foster, 1970; Hays et al., 1969; Opdyke and Glass, 1969; Hunkins et al., 1971; Goodell and Watkins, 1968). The maximum age of the sediments in these studies was limited by the penetration of piston coring devices. The longest conventional piston core in the LDGO collection is about 26 m and, therefore, a stratigraphic record to the basal Pliocene was extremely rare. Cores recording most of the PlioPleistocene magnetic polarity record included E13-17 (Hays and Opdyke, 1967 ), RC 12-66 (Foster and Opdyke, 1970) (Fig. 6.2), and V28-179 (Fig. 6.1) (Shackleton and Opdyke, 1977). In early studies, the durations of the Plio-Pleistocene subchrons were estimated through the magnetic stratigraphy of deep-sea cores by extrapolation of sedimentation rates from reversal transitions or by interpolation assuming constant sedimentation rates and adopting K-Ar ages for polarity chron boundaries (Fig. 6.3). Estimates were compared with those obtained from high-resolution oceanic magnetic anomaly profiles assuming constant sea floor spreading rates. The last decade has witnessed a dramatic improvement in the quantity and quality of the Plio-Pleistocene magnetostratigraphic records due, in large part, to the advent in the early 1980s of the hydraulic piston corer (HPC) and, subsequently, the advanced piston corer (APC). The HPC, first used during Leg 68 of the Deep Sea Drilling Project (DSDP), allowed recovery of several hundred meters of sediment by piston coring, increasing the penetration of piston coring by more than an order of magnitude. The HPC was a considerable improvement over the standard rotary drilling techniques, which often resulted in poor core recovery and drilling-related deformation. Notable among the Plio-Pleistocene magnetostratigraphies recovered by HPC/APC are those from DSDP Leg 68 in the Carribbean (Kent and Spariosu, 1983), DSDP Leg 73 in the South Atlantic (Tauxe et al., 1983c, 1984; Poore et al., 1984), and DSDP Leg 94 in the North Atlantic (Weaver and Clement, 1986). Apart from their value for biomagnetostrati-

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graphic correlations, the Plio-Pleistocene magnetic stratigraphies from DSDP Leg 94 were important for two additional reasons. First, as mentioned above, they resulted in the documentation of two short-duration normal polarity subchrons within the Matuyama chron (Gilsa and Cobb Mountain subchrons) which had not previously been adequately recorded in sediments (Fig. 6.1) (Clement and Kent, 1987). Second, the correlation of oxygen isotope records to Leg 94 Plio-Pleistocene magnetic stratigraphies resulted in a first attempt to astronomically tune this part of the GPTS (Ruddiman et al., 1986, 1989; Raymo et al., 1989). The vast majority of Plio-Pleistocene marine magnetostratigraphic records are from pelagic deep-sea sediments. Plio-Pleistocene shallow water platform carbonates recovered from boreholes in the Bahamas (McNeill et al., 1988) and Mururoa Atoll (Aissaoui et al., 1990) have yielded excellent

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6 The Pliocene-Pleistocene Polarity Record

magnetic stratigraphies despite pervasive dolomitization. These magnetostratigraphies provide a powerful chronostratigraphic tool in an environment where biostratigraphic control is inherently poor. It is important to balance these important magnetostratigraphic studies in Plio-Pleistocene shallow water limestones with the many unsuccessful (and usually unpublished) magnetostratigraphic studies in ancient shallow water limestones which are generally thwarted by remagnetization or very weak remanence intensities. The Plio-Pleistocene stages in most geological time scales are derived from stratotype sections located in Italy (see Rio et al., 1991). Correlation of these stratotype sections to marine sequences outside the Mediterranean is complicated by provinciality of Mediterranean flora and fauna and by changes in lithofacies in the stratotype sections. Magnetic stratigraphies in the stratotype region provide the potential means of global correlation. The focus of Early Pliocene Mediterranean magnetostratigraphies has been the Trubi Limestones, a fine-grained pelagic limestone recording the postMessinian flooding of the Mediterranean and cropping out in Sicily and Calabria (southern Italy). Magnetostratigraphic records of the MiocenePliocene boundary and the Gilbert Chron in Calabria (Zijderveld et al., 1986; Channell et al., 1988) were expanded by the development of a composite reference section in Sicily (Rossello composite) from the MiocenePliocene boundary to the base of the Matuyama Chron (Fig. 6.4) (Langereis and Hilgen, 1991). The Trubi Limestones are overlain by a more marly facies (Monte Narbone Formation), the transition occurring in the middle part of the Gauss Chron. Upward extension of magnetic stratigraphies into the Late Pliocene and Early Pleistocene has been documented in Calabria (Zijderveld et al., 1991) and Sicily (Zachariasse et al., 1989, 1990). The Plio-Pleistocene boundary stratotype at Vrica (Calabria) originally studied magnetostratigraphically by Tauxe et al. (1983a) has been restudied by Zijderveld et al. (1991), and the stage boundary placed close to, but just above, the top of the Olduvai subchron. The land section magnetostratigraphic data in southern Italy are far better quality than those obtained from rotary cores during ODP Leg 107 in the Tyrrhenian Sea (Channell et al., 1990b), although biostratigraphic control is better in the Tyrrhenian cores due to enhanced nannofossil preservation (Rio et al., 1990). The biomagnetostratigraphic correlations which have resulted from land section studies in Italy and ODP Leg 107 cores (Fig. 6.5) facilitate export of PlioPleistocene stages from the Mediterranean and are an important basis for global Plio-Pleistocene correlations. Although the Early Pliocene (Zanclean) type section is within the Sicilian Trubi Limestones, other Pliocene stratotype sections (Tabianian and Piacenzian) are in the terrigeneous "argille azzurre" (blue clay) facies

6.3 Magnetic Stratigraphy in Plio-Pleistocene Sediments

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104

6 The Pliocene-Pleistocene Polarity Record

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6.3 Magnetic Stratigraphy in Plio-Pleistocene Sediments

]05

of the Northern Apennines (see Rio et al., 1991). This is a more problematic facies for magnetic stratigraphy due to the ubiquitous presence of iron sulfides, although Pliocene magnetic stratigraphies are available from a few sections in this region (Mary et al., 1993; Channell et al., 1994). The available magnetic stratigraphies indicate that sedimentation in the Northern Apennine foredeep in the vicinity of the stratotype sections was very discontinuous and that the region is therefore not suitable for stratotype designations. The Miocene-Pliocene boundary stratotype is a lithologic boundary in the stratigraphic sequence at Capo Rosello (Sicily) (Fig. 6.4). Benson and Hodell (1994) argue that it would be better to redefine the Miocene/ Pliocene boundary and correlate it to the base of the Gilbert Chron in Morocco, because of the inadequate biostratigraphic definition of the boundary in Sicily. The Pliocene-Pleistocene boundary at the stratotype section (Vrica, Italy) is defined by the first occurrence of migrant taxa such as A r t i c a islandica which cannot be easily correlated to the open ocean. These species first appear in the Vrica section just above the Olduvai subchron (Tauxe et al., 1983a; Zijderveld et al., 1991). A definition of the Pliocene-Pleistocene boundary at the top of the Olduvai subchron should perhaps be reconsidered. Berggren et al. (1980) gave 37 biostratigraphic datums for the PlioPleistocene of the world ocean and many more have subsequently been added (Fig. 6.6). The apparent correlation of bioevents to the GPTS can vary from site to site due to disconformities, fossil preservation, and the operational method used to define the bioevent and the species itself. Many of these datums have been correlated directly to the 61sO time scale. Emiliani (1955) produced the first long records of the change in 61sO in the oceans through time in foraminiferal tests relative to a standard (a belemnite from the Pee Dee Formation in South Carolina). The ocean water was shown to oscillate between times when the ocean had 61sO ratios similar to those today (interglacial conditions) to times when the oceans were enriched in 180 (glacial conditions). Emiliani numbered these changes beginning at the present interglacial (stage 1) and the last glacial (stage 2). In this scheme, the interglacials are indicated by odd numbers and the glacial stages by even numbers. Shackleton and Opdyke (1973) extended this nomenclature to below the Brunhes/Matuyama boundary and correlated this boundary to isotopic stage 19. The 61sO record was extended to the entire Pleistocene by Shackleton and Opdyke (1976) and to the base of the Matuyama by Ruddiman et al. (1986), who extended the numbering system to isotopic stage 102 at the Gauss/Matuyama boundary. The isotopic variations since ---1 Ma were shown to have a dominant periodicity of ---100 ky (Hays et al., 1976). For the time from the Late Jaramillo to the base of the Matuyama, the variations in 61sO have been shown to have a dominant

106

6 The Pliocene-Pleistocene Polarity Record

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Figure 6.6 The correlation to the Plio-Pleistocene GPTS of first and last appearances of selected fauna and flora. F A D = first appearance datum, L A D = last appearance datum.

6.4 Astrochronologic Calibration of the Plio-Pleistocene GPTS

107

periodicity of ---42 ky (Ruddiman et aL, 1986). The isotopic sequence has been extended by Shackleton et al. (1995b) to the base of the Gilbert Chron at Site 846 from the eastern Pacific. The sequential numbering system was discontinued below the Gauss/Matuyama boundary and a letter prefix was used to label isotopic stages in the Gauss and Gilbert Chrons (Fig. 6.7).

6.4 Astrochronologic Calibration of the Plio-Pleistocene GPTS Hays et al. (1976) were the first to demonstrate the presence of orbital cycles of eccentricity, obliquity, and precession in paleoclimate proxies in deep-sea sediments. These authors adjusted their initial time scale to bring the peak variance of the obliquity cycles to the same frequency calculated for obliquity by astronomers, and were thus the first to use cyclostratigraphy as a means of tuning time scales. During the 1980s, astrochronology was largely restricted to the Brunhes Chron and was utilized successfully to constrain the age of the oxygen isotope record with a precision of a few thousand years (e.g., Martinson et al., 1987). The turning point for astrochronology/cyclostratigraphy came with DSDP/ODP hydraulic piston core recovery at multiple holes for single sites with adequate sedimentation rates, oxygen isotopic records, and magnetic stratigraphies. The hydraulic piston corer allowed good undisturbed sediment core recovery, and the drilling of multiple holes allowed complete recovery of composite sections. Ruddiman et al. (1989) and Raymo et al. (1989) compiled oxygen isotope data for the Pleistocene and Upper Pliocene at DSDP Leg 94 Site 607 (North Atlantic). Orbital tuning using the strong obliquity signal for the Matuyama Chron did not reveal any significant discrepancy with the standard (Mankinen and Dalrymple, 1979) K-Ar polarity chron ages, apart from a significantly shorter duration for the Olduvai subchron. For over ten years, the Mankinen and Dalrymple (1979) compilation of K-Ar ages of Plio-Pleistocene polarity chrons was generally accepted in spite of astrochronology indicating a significantly older age for the Brunhes/ Matuyama boundary (Johnson, 1982). The realization that the Mankinen and Dalrymple (1979) polarity chron ages may be in error came with the study of ODP Site 677 in the eastern equatorial Pacific (Shackleton et al., 1990). At this site, the oxygen isotopic record from benthic foraminifera is dominated by the obliquity signal and the planktic record by precession. Because the effect of precession is modulated by the eccentricity cycle, the precession-dominated planktic record provided a better means of correlation to the astronomical data than the obliquity record. The Site 677 oxygen isotope data could therefore be matched to the astronomical data with more confidence than the poorly modulated obliquity signal at Site

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6 The Pliocene-Pleistocene Polarity Record

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et

al., 1990, 1995b).

607. Indeed at Site 607, Ruddiman et al. (1989) apparently miscounted the number of obliquity cycles in the lower Brunhes and early Pleistocene partly because they accepted the K-Ar age for the Brunhes/Matuyama boundary. The Site 677 astrochronology (Shackleton et al., 1990) yielded an age for the Brunhes/Matuyama boundary (0.78 Ma) at odds with the K-Ar age (0.73 Ma). The older astrochronological age for the Brunhes/Matuyama boundary from Site 677 coincided with the observation of discrepancies between

6.4 AstrochronologicCalibration of the Plio-Pleistocene GPTS

|09

conventional K-Ar ages and Pliocene astrochronology derived from carbonate cyclostratigraphy in the southern Italian Trubi Marls (Hilgen and Langereis, 1989). Spectral analyis of CaCO3 content in these Pliocene carbonates yielded periodicities of 15.5, 18.5, 35.0, and 335 ky, indicating a consistent discrepancy with the astronomical solutions of 19, 23, 41, and 413 ky. Hilgen and Langereis (1989) computed the age of the Gilbert and Gauss reversals using an age of 3.40 Ma for the Gauss/Matuyama boundary and tuning the CaCO3 cycles to the astronomical solution. The Late Pliocene and Early Pleistocene of southern Italy is characterized by sapropels, dark organic-rich laminated layers. Hilgen (1991a) showed that for Late Pleistocene sapropels from the eastern Mediterranean (core RC 9-181), individual sapropels correlate to minimum peak values of the precession index and small- and large-scale sapropel clusters correlate to eccentricity maxima related to the 100 ky and 400 ky eccentricity cycles. Hilgen (1991a) applied this observation to the sapropel-bearing Late Pliocene-Early Pleistocene sections in southern Italy which had magnetostratigraphic age control. The sapropel occurrences and sapropel clusters could be matched to the astronomical solution leading to a new time scale for the 1.8-3.2 Ma interval. Hilgen (1991a) illustrated how the time scales of Berggren et al. (1985b) and Raymo et al. (1989) lead to a mismatch of the sapropel occurrences with the astronomical solution. Below the main interval of sapropel occurrence, Hilgen (1991b) used the CaCO3 cycles in the same southern Italian sections to generate an astrochronology for the Pliocene. Hilgen (1991b) used the relationship between the sapropel occurrences and CaCO3 content in the Upper Pliocene, to infer that the gray marl beds denoting the small-scale CaCO3 minima correspond to minima in the precession index and that the larger scale CaCO3 minima correspond to maximum amplitude variations of the precession index related to the eccentricity modulation. The recognition of both a precession and an eccentricity signal in the CaCO3 record allowed the observations to be tuned to the astronomical solution with considerably more confidence than would have been the case in the absence of the eccentricity modulation. This important paper resulted in astrochronological ages for Gauss and Gilbert reversals (Table 6.2). Neogene astrochronology has been extended into the Late Miocene using cores collected during ODP Leg 138 (Shackleton et al., 1992, 1995a). Multiple cores at each site allowed composite sections to be compiled at each site using magnetic susceptibility and GRAPE (Gamma Ray Attenuation Porosity Evaluator) data (Hagelberg et al., 1992). The GRAPE records reflect the ratio of calcite to biogenic opal, and the GRAPE records from the composite sections were matched with the orbital insolation solution of Berger and Loutre (1991). This matching has led to astrochronological

~

~

Table 6.2 Pliocene-Pleistocene Time Scales E

Schackleton el al. [1995a] (Ma) tr

0.78 0.99 1.07 1.77 1.95 2.60

Cande and Kent [1992a] (Ma)

"

0.73 0.91 0.98 1.66 1.88 2.47 2.92 2.99 3.08 3.18 3.40 3.88 3.97 4.10 4.24 4.40 4.47 4.57 4.77 5.35

Hilgen [1991a.b] (Ma) ~

Shackleton er a/. [I9901 (Ma) ..,.d

t~

Berggren et a/. [1985] (Ma) tr

Mankinen and Dalrymple [1979] (Ma) ,...,

O

Polarity chron boundary

Baksi [1994] (Ma)

Cande and Kent [1995]

0.78 0.99 1.05 1.78 2.02 2.64 3.10 3.17 3.27 3.38 3.61 4.12 4.23 4.37 4.55 4.74 4.82 4.94

0.78 0.99 1.07 1.77 1.95 2.58 3.04 3.11 3.22 3.33 3.58 4.18 4.29 4.48 4.62 4.80 4.89 4.98 5.23 5.89

~~

~ I~

,,-

~'3

r

tr

e,i

~

~ ,--~ ~,1 ~'3 tr ,.--~ ~',i ',~" ~,I~ OO OO ~'~ ~,100

~ ~'~

3.046 3.131 3.233 3.331 3.594 4.199 4.316 4.479 4.623 4.781 4.878 4.977 5.232 5.882

~

~

~'~

2.5712.62 3.04 3.11 3.22 3.33 3.58 4.18 4.29 4.48 4.62 4.80 4.89 4.98 5.23

0.780 0.984 1.049 1.757 1.983 2.600 3.054 3.127 3.221 3.325 3.553 4.033 4.134 4.265 4.432 4.61 1 4.694 4.812 5.046 5.705

t"-

c5

t~ t~ ~ :~ :~

0.73 0.90 0.97 1.67 1.87 2.48 2.92 3.01 3.05 3.15 3.40 3.80 3.90 4.05 4.20 4.32 4.47 4.85 5.00

c5

E

Brunhes/Matuyama Jaramillo (top) Jaramillo (base) Olduvai (t) Olduvai (b) Matuyama/Gauss Kaena (t) Kaena (b) Mammoth (t) Mammoth (b) GausslCilbert Cochiti (t) Cochiti (b) Nunivak (t) Nunivak (b) Sidufjall (t) Sidufjall (b) Thvera (t) Thvera (b) Gilbert (b)

6.5 4~

Age Calibration of the Plio-Pleistocene GPTS

|||

age estimates for geomagnetic reversals in the 3-6.2 Ma (top Kaena to C3A.n2) interval. Consistent estimates were derived from several holes/ sites. The estimates are within a few tens of thousands of years of the estimates given by Hilgen (1991a,b) from southern Italian sections and differ from those given by Cande and Kent (1992a) by up to several hundred thousand years. Finally, these authors (Shackleton et al., 1995a) generated a sea floor anomaly time scale beyond C3A.2n by utilizing their age for the top of C3A.ln (5.875 Ma), the radiometric age of Baksi (1993) for the top of C5n.ln (9.639 Ma), and the sea floor distances given by Cande and Kent (1992a). How far back in time the astrochronological technique can be carried will rest on the availability of a suitable proxy climatic records in the oceans or cropping out on land (see Hilgen et al., 1995).

6.5 4~

Age Calibration of the Plio-Pleistocene GPTS

The advent of high-precision 4~ age dating techniques and the implication from cyclostratigraphy that the conventional K-Ar ages for the PlioPleistocene (e.g., Mankinen and Dalrymple, 1979) were too young have studies aimed at testing the astroresulted in a large number of 4~ chronological ages of Plio-Pleistocene polarity chrons. The conventional K-Ar age for the Brunhes/Matuyama boundary (0.73 Ma) has been modified studies (Izett and to 0.78 Ma by a number of independent 4~ Obradovich, 1991; Baksi et al., 1992; Spell and McDougall, 1992; Tauxe et al., 1992; Hall and Farrell, 1993). This age is consistent with the cyclostratigraphic estimates giving strong support to the new chronology (recent 4~ 39Ar ages summarized by Baksi, 1994, see Table 6.2). The astrochronological ages derived by Shackleton et al. (1990) and Hilgen (1991a,b) have now ages for the Jaramillo (Glass et al., 1991; Spell been supported by 4~ and McDougall, 1992; Tauxe et al., 1992), Cobb Mountain (Turrin et al., 1994), R6union (Baksi et al., 1993b), Olduvai (Walter et al., 1991; Baksi, 1994). Kaena and Mammoth (Renn6 et al., 1993; Walter, 1994; Walter and Aronson, 1993), and Gilbert (McDougall et al., 1992; Baksi et al., 1993a) (Fig. 6.8). The 4~ ages of Plio-Pleistocene polarity chrons (reviewed by Baksi, 1994) have in general confirmed the astrochronological determinations, thereby ratifying the astrochronology (Table 6.2). The ages for PlioPleistocene reversals based on oceanic magnetic anomalies given by Cande or astrochronological and Kent (1992a) are not consistent with the 4~ estimates, particularly for the Gilbert Chron (Table 6.2). D. S. Wilson (1993) analyzed the magnetic anomalies at various spreading centers and concluded that the anomaly spacing is, in fact, consistent with the astrochronological estimates. Cande and Kent (1992a) adopted the spacing of Plio-

l 12 ~:E

6 The Pliocene-Pleistocene Polarity Record 4

t~

mO

O tO

I,_ c-

3

8

t.o 03

E o t..

2

00

t~ 03 t-

O .m ta3 t._

03 O 12.

0

0

,

I

1

,

I

2

,

I

3

Polarity transition K/At and Ar/Ar ages (Ma) Figure 6.8 Comparison of geomagnetic polarity reversal ages based on orbital tuning (Shacklefeldspar ages ton et al., 1990; Hilgen 1991a,b) with reversal ages based on K/Ar and 4~ from the Turkana Basin (Spell and McDougall, 1992; McDougall et al., 1992).

Pleistocene anomalies given by Klitgord et al. (1975), which appears to be the source of the error. In Cande and Kent (1995), the astrochronological estimates for the age of the Plio-Pleistocene reversals were adopted (Table 6.2). The consistency of the astronchronological, 4~ and sea floor age estimates for Plio-Pleistocene anomalies has ratified the astrochronological techniques. The resolution of astrochronological estimates for the age of reversal boundaries in the Plio-Pleistocene has reached one precession cycle (---20 ky), only a factor of two greater than the time taken for field reversal to occur. Renn6 et al. (1994) have taken the next step and used the astrochronological estimates to calibrate the 4~ dating standard (Fish Canyon sanidine, Mmhb-1), thereby reducing the uncertainty to 0.6% for 4~ ages calibrated against this standard.

Late Cretaceous-Cenozoic GPTS

7.1 Oceanic Magnetic Anomaly Record Vine and Matthews (1963) proposed that lineated magnetic anomalies observed at sea using towed magnetometers resulted from the permanent magnetization of the oceanic crust. They suggested that this magnetic signature was a result of normal and reverse polarity geomagnetic fields and the sea floor spreading process outlined by Hess (1962). This hypothesis was essentially confirmed by Vine and Wilson (1965), Pitman and Heirtzler (1966), and Vine (1966). These papers demonstrated that: (1) the magnetic anomaly pattern was generally symmetric on both sides of a spreading ridge, and (2) the anomaly pattern emanating from the ridge reproduces in detail the reversal sequence obtained by the radiometric dating of lava flows on land. As a result, Vine (1966) and Pitman and Heirtzler (1966) extended the magnetic reversal pattern to 10 Ma and dated it by assuming constant spreading rate and an age for the base of the Gauss Chron. It rapidly became apparent that correlatable magnetic anomalies generated by sea floor spreading extend for thousands of km away from the ridge crests of the North Pacific, South Pacific, and South Atlantic oceans (Fig. 7.1). This led to the realization that it might be possible to extend the geomagnetic polarity time scale (GPTS), as derived from these anomalies, back in time to the late Cretaceous with reasonable accuracy if the spreading rates remained constant. Heirtzler et al. (1968) compared anomaly patterns from all oceans. Taking into consideration what was known about the age of the sea floor at that time, they took a courageous step by deciding that the South Atlantic Ocean had spread at an almost constant rate since the Late 113

| 14

7 Late Cretaceous-Cenozoic GPTS

u

t13

A

~o

t...

v .<

-o--

-o

F

I

i

~

=

i.-~ J~ m

-_...

I" o

i I

s.

in N

p~o

.,9

E

E

&

,

E &

w

.J

o

o Ii

Figure 7.1

E &

fBO

mm

E &

~

E

&

E

,,

u_ d

"i"

>

o= z

2,

w

Oceanic magnetic anomaly profiles from South Atlantic (V-20), North Pacific (V16), and South Pacific (SI-6, EL-19S) with synthetic model profiles. Time scale constructed assuming constant seafloor spreading rates and an age of 3.35 Ma for the base of the Gauss Chron (after Heirtzler et al., 1968).

7.1 Oceanic Magnetic Anomaly Record

| |5

Cretaceous. They then simulated the polarity pattern necessary to give the observed pattern of magnetic anomalies, assumed an age of 3.35 Ma for the base of the Gauss, and assigned ages by extrapolation assuming a constant spreading rate. With this bold and imaginative step, Heirtzler et al. (1968) produced a dated polarity pattern for the entire Cenozoic and part of the Cretaceous and deduced an age of 60 Ma for the K/T boundary. This geomagnetic polarity time scale (GPTS) became a vital tool for deciphering the age and tectonic history of the ocean basins. In their original treatment of magnetic anomalies, Heirtzler et al. (1968) assigned numbers to certain prominent magnetic anomalies, from anomaly 1 (Brunhes Chron) at the ridge crest to anomaly 32 (Late Cretaceous) (Fig. 7.1). The labeling of magnetic anomalies was found to be very useful as an aid to correlation. As more detailed surveys became available, however, it became clear that a finer division of the anomaly pattern was desirable. A study by Blakely (1974) of closely spaced magnetic anomalies in the Neogene Pacific led to the adoption of a number/letter scheme for magnetic anomalies in the Miocene, a change made necessary by the high reversal rate during that time. Since the publication of Heirtzler et al. (1968), most published geomagnetic polarity time scales (e.g., LaBrecque et al., 1977; Berggren et al., 1985a,b; Harland et al., 1990) utilized the basic Heirtzler et al. (1968) anomaly sequence interpretation (Fig. 7.2), albeit with some important modifications for anomalies 4A to 6 (Blakely, 1974), anomalies 1 to 3A (Klitgord et al., 1975), and anomalies 30 to 34 (Cande and Kristoffersen, 1977). The most recent and extensive revision of the geomagnetic polarity time scale for the Late Cretaceous and Cenozoic has been made by Cande and Kent (1992a, 1995). In most recently derived Late Cretaceous-Cenozoic geologic time scales, absolute ages are correlated to the GPTS, which is then utilized to interpolate between absolute age estimates. The template for the GPTS is derived from oceanic magnetic anomalies using the constant spreading rate assumption. The absolute ages of geologic stage boundaries can then be estimated if the stage boundaries are correlated to the GPTS. In this way, the GPTS is central to the construction of geologic time scales and provides the means to correlate among the diverse measures of geologic time, such as biostratigraphy, isotope stratigraphy, and absolute ages. In some time scales, the link between geologic stages and absolute ages has been attempted without use of the GPTS as a means of interpolation (e.g., Odin et al., 1982); however, it is now generally accepted that the GPTS should be the central thread of Late Cretaceous-Cenozoic time scales. Cande and Kent (1992a), after surveying oceanic magnetic anomalies from the world's oceans, concluded, as had Heirtzler et al. (1968), that the

I ]6

7 Late Cretaceous-Cenozoic GPTS

-r

,.. t--

_J

o

-

|

.J 13,. 0

,,~

0Mo

r-,. r,,-

~

o ~

O)

,-~,

//"

ao >

ff<_fi S N '

I

~ ~

--

i

"

/

-

~ QO

o) r

r

I~

o oo "

~

w

-45

5 ~:

10

'

- - 5 0

/

I

~

-

-55

15

/

,

-

/

"

_

/

"

-

,

-65

,

- 7 0

,

-

-60Ma

20Ma

25

:30

-75 35 -80Ma

40Ma

y -85

45 -90

7.1 Oceanic Magnetic Anomaly Record

1 17

South Atlantic magnetic anomalies were the most appropriate as the basis for a new GPTS. The South Atlantic has a well studied set of magnetic anomalies (Cande e t al., 1988) to anomaly 34. The South Atlantic also has both flanks of the ridge crest preserved and has not suffered from major plate reorganization. It does, however, have a rather slow rate of spreading. The procedure of Cande and Kent (1992a) was to use a set of nine finite rotation poles and determine the distance to nine conjugate magnetic anomalies. These were positioned on a flow line at approximately 30~ latitude. The chosen anomalies (4A, 5C, 7, 13, 20, 24, 30, 33, and 34) were positioned at intervals of 150 to 300 km along the flow line. These segments were designated category 1 intervals. The details of the anomaly pattern between these category 1 intervals were filled in by choosing undisturbed segments of anomaly sequences from individual profiles elsewhere in the South Atlantic. These sequences were then continued downward 1.5 to 2 km, deskewed, and the crossover points on the profiles were then used to indicate the position of the reversal boundaries on the sea floor record (Fig. 7.3). Five to nine individual records were stacked after adjustment to the previously determined tie points (category 1 intervals) (e.g., Fig. 7.4). The positions and widths of the anomalies were then determined by averaging. This process yielded an average error of about 7%. In order to obtain the finer detail of the anomaly sequence, known from fast spreading ridges in the Pacific and Indian oceans, individual studies of anomaly sequences from these oceans were inserted between the tie points. For instance, the studies of Klitgord e t al. (1975) and Blakely (1974) (Fig. 7.5) were used to determine the fine-scale reversal sequence between the ridge crest and anomaly 6. Very low amplitude, but correlatable, magnetic anomalies have been called "tiny wiggles" (Cande and LaBrecque, 1974; Blakely, 1974) which can be modeled as short polarity chrons or as intensity fluctuations of the geomagnetic field. Some have already been identified as short polarity chrons, such as the Rdunion subchron (Gromm6 and Hay, 1971) and the Cobb Mountain subchron (Mankinen e t a L , 1978, Mankinen and Gromm6, 1982; Clement and Kent, 1987), which have durations of about 20 ky. Cande and Kent (1992b) have suggested that many of the correlatable tiny wiggles

Figure 7.2 Evolution of Late Cretaceous-Cenozoic geomagnetic polarity time scales from Hiertzler et al. (1968) (HDHPL68) to Cande and Kent (1995) (CK95) (after Mead, 1996). Key: HDHPL68: Heirtzler et al. (1968),TM76: Tarling and Mitchell (1976), LKC77: LaBrecque et al. (1977), MD79: Mankinen and Dalrymple (1979), NLC80: Ness et al. (1981), LA81: Lowrie and Alvarez (1981), GTS82: Harland et al. (1982), BKFV85: Berggren et al. (1985a,b), HHV88: Haq et al. (1987), GTS89: Harland et al. (1990), CK92: Cande and Kent (1992a), CK95: Cande and Kent (1995).

1 18

7 Late Cretaceous-Cenozoic GPTS

27

2829

30 31

32

REVERSAL BOUNDARIES

uuV

DESKEWED

W

DOWNWARD CONTINUED

ORIGINAL PROFILE Figure 7.3 Procedure for constructing block models from oceanic magnetic anomaly profiles. Original profile is projected perpendicular to the strike of the lineations, bandpassed, downward continued, deskewed, and reversal boundaries determined from zero crossings (after Cande and Kent, 1992a).

AXIS

,

2

2A

I,., / ' ~

~v,V

3

3A

1~,~

,~--

V

4

If'-,',,

,,~

I/"',,

i STAC K

,,,> m

~600nT

25km t

, I

kt\/

"-'~

\ / - -

I~

k,.,4

V

L.,/

',..4

--'1

-,~

kJ I

.,,,j~

A2

(~1 k AA I . Lc', _ 1 ~ I ~k,,,../l~/1 " ~ ' q x ' ~ k ~ " VN'- - ~ - " A3 I

ivx.,4 ~1 \ #

I qv~ ,

IA A

"k,/I - ',J"1

"t/~

W jvV

I/-'k

I An

I A6

v

Figure 7.4 South Atlantic stack for anomaly 1 to anomaly 4. Distances in the stack (category 2 distances) are determined by averaging the widths of the subintervals on the individual profiles (after Cande and Kent, 1992a).

7.2 Numerical Age Control

1

J

2

2A

I]

3

_NI

Wilson & Hey, 1981 Rea and Blakeley, 1975

]

~

[-1__~__~_]~]~

[

]

l

[

I

Klitgord et al. 1975 ] SouthAtlantic (category 2 distances)

i

i

20 km Figure 7.5 Block models for anomaly 1 to anomaly 3A from fast-spreading centers integrated within the South Atlantic (category 2) distance intervals (after Cande and Kent, 1992a).

are related to geomagnetic intensity variations and have coined the term "cryptochron" for these short duration intervals.

7.2 NumericalAge Control The Deep Sea Drilling Project, which began in 1968, provided a wealth of biostratigraphic data for sediments directly overlying magnetic anomalies, providing a minimum paleontological age for many anomalies. Improved correlations of magnetic anomalies to the geologic time scale have subsequently been achieved by correlations of magnetic anomalies to polarity zones in magnetostratigraphic sections with good biostratigraphic control. It is generally not possible to obtain radiometric ages on the sea floor lavas themselves, since they have suffered extensive alteration and argon does not completely outgas at deep ocean floor pressures. Absolute ages are generally correlated to magnetic anomaly records by a two-step process: correlation of magnetic anomalies to magnetostratigraphic sections and correlation of absolute ages to magnetostratigraphic sections. The Cande and Kent (1992a) time scale (Table 7.1) utilized the following nine absolute age tie points for the last 83 Ma. (1) Gauss/Matuyama boundary at 2.60 Ma from astrochronology (Shackleton et al., 1990). (2) C5Bn at 14.8 Ma through the correlation of foraminiferal zonal boundary N9/N10 to the GPTS (Miller et al., 1985; Berggren et al., 1985b) and the age estimates for this foraminiferal boundary from Japan at 14.5 _+ 0.4 Ma (Tsuchi et al., 1981) and from Martinique at 15.0 _+ 0.3 Ma (Andreieff et al., 1976). (3) Oligocene-Miocene boundary correlated to the middle part of C6Cn (Berggren et al., 1985b), chronogram ages for this stage boundary (23.8 Ma) from Harland et al. (1990). (4) EoceneOligocene boundary correlated to chron C13r (.14) in the Apennines (Nocchi et al., 1986), with absolute age estimate (33.7 _+ 0.4 Ma) from

120

7 Late Cretaceous-Cenozoic GPTS

Table 7.1 Normal Polarity Intervals (Ma) Polarity chron

Cande and Kent (1992a)

Cande and Kent (1995)

Wei (1995)

Cln Clr.ln C2n C2r.ln C2An.ln C2An.2n C2An.3n C3n.ln C3n.2n C3n.3n C3n.4n C3An.ln C3An.2n C3Bn C3Br.ln C3Br.2n C4n.ln C4n.2n C4r.ln C4An C4Ar.ln C4Ar.2n C5n.ln C5n.2n C5r.ln C5r.2n C5An.ln C5An.2n C5Ar.ln C5Ar.2n C5AAn C5ABn C5ACn C5ADn C5Bn.ln C5Bn.2n C5Cn.ln C5Cn.2n C5Cn.3n CSDn C5En C6n C6An.ln C6An.2n C6AAn C6AAr.ln

0.000-0.780 0.984-1.049 1.757-1.983 2.197-2.229 2.600-3.054 3.127-3.221 3.325-3.553 4.033-4.134 4.265-4.432 4.611-4.694 4.812-5.046 5.705-5.946 6.078-6.376 6.744-6.901 6.946-6.981 7.153-7.187 7.245-7.376 7.464-7.892 8.047-8.079 8.529-8.861 9.069-9.146 9.428-9.491 9.592-9.735 9.777-10.834 10.940-10.989 11.378-11.434 11.852-12.000 12.108-12.333 12.618-12.649 12.718-12.764 12.941-13.094 13.263-13.476 13.674-14.059 14.164-14.608 14.800-14.890 15.038-15.162 16.035-16.318 16.352-16.515 16.583-16.755 17.310-17.650 18.317-18.817 19.083-20.162 20.546-20.752 21.021-21.343 21.787-21.877 22.166-22.263

0.000-0.780 0.990-1.070 1.770-1.950 2.140-2.150 2.581-3.040 3.110-3.220 3.330-3.580 4.180-4.290 4.480-4.620 4.800-4.890 4.980-5.230 5.894-6.137 6.269-6.567 6.935-7.091 7.135-7.170 7.341-7.375 7.432-7.562 7.650-8.072 8.225-8.257 8.699-9.025 9.230-9.308 9.580-9.642 9.740-9.880 9.920-10.949 11.052-11.099 11.476-11.531 11.935-12.078 12.184-12.401 12.678-12.708 12.775-12.819 12.991-13.139 13.302-13.510 13.703-14.076 14.178-14.612 14.800-14.888 15.034-15.155 16.014-16.293 16.327-16.488 16.556-16.726 17.277-17.615 18.281-18.781 19.048-20.131 20.518-20.725 20.996-21.320 21.768-21.859 22.151-22.248

0.000-0.780 0.990-1.070 1.770-1.950 2.140-2.150 2.580-3.040 3.110-3.220 3.330-3.580 4.180-4.290 4.480-4.620 4.800-4.890 4.980-5.230 5.829-6.051 6.173-6.450 6.795-6.943 6.986-7.019 7.183-7.216 7.271-7.398 7.483-7.902 8.055-8.088 8.543-8.887 9.106-9.191 9.490-9.560 9.670-9.827 9.874-11.089 11.214-11.273 11.738-11.806 12.314-12.495 12.628-12.904 13.256-13.294 13.379-13.436 13.654-13.843 14.050-14.312 14.555-15.021 15.147-15.677 15.901-16.007 16.178-16.320 17.289-17.592 17.628-17.800 17.872-18.051 18.617-18.955 19.603-20.074 20.321-21.295 21.633-21.814 22.047-22.324 22.703-22.780 23.025-23.107

7.2 Numerical Age Control

Table 7.1

| 2 |

continued

Polarity chron

Cande and Kent (1992a)

Cande and Kent (1995)

Wei (1995)

C6AAr.2n C6Bn.ln C6Bn.2n C6C n. 1n C6Cn.2n C6Cn.3n C7n.ln C7 n.2n C7An C8n. 1n C8n.2n C9n C 1On. 1n C10n.2n C1 l n . l n C11n.2n C12n C13n C15n C16.1n C 16n.2n C17n.1n C17n.2n C 17n.3n C18n.ln C18n.2n C19n C20n C21 n C22n C23n. 1n C23n.2n C24n. 1n C24n.2n C24n.3n C25n C26n C27n C28n C29n C30n C31n C32n.ln C32n.2n C32r. 1n C33n C34n

22.471-22.505 22.599-22.760 22.814-23.076 23.357- 23.537 23.678-23.800 23.997-24.115 24.722-24.772 24.826-25.171 25.482-25.633 25.807-25.934 25.974-26.533 27.004-27.946 28.255- 28.484 28.550-28.716 29.373-29.633 29.737-30.071 30.452-30.915 33.050-33.543 34.669-34.959 35.368-35.554 35.716- 36.383 36.665-37.534 37.667-37.915 37.988-38.183 38.500-39.639 39.718-40.221 41.353-41.617 42.629-43.868 46.284-47.861 48.947-49.603 50.646-50.812 50.913-51.609 52.238- 52.544 52.641-52.685 52.791-53.250 55.981-56.515 57.800-58.197 61.555-61.951 63.303-64.542 64.911-65.732 66.601-68.625 68.745-69.683 71.722-71.943 72.147-73.288 73.517-73.584 73.781-78.781 83.000-

22.459-22.493 22.588-22.750 22.804-23.069 23.353- 23.535 23.677-23.800 23.999-24.118 24.730-24.781 24.835-25.183 25.496-25.648 25.823-25.951 25.992-26.554 27.027-27.972 28.283- 28.512 28.578-28.745 29.401-29.662 29.765-30.098 30.479-30.939 33.058-33.545 34.655-34.940 35.343-35.526 35.685- 36.341 36.618-37.473 37.604-37.848 37.920- 38.113 38.426-39.552 39.631-40.130 41.257-41.521 42.536-43.789 46.264-47.906 49.037-49.714 50.778-50.946 51.047-51.743 52.364- 52.663 52.757-52.801 52.903-53.347 55.904-56.391 57.554-57.911 60.920-61.276 62.499-63.634 63.976-64.745 65.578-67.610 67.735-68.737 71.071-71.338 71.587-73.004 73.291-73.374 73.619-79.075 83.000-

23.284-23.312 23.392-23.527 23.572-23.794 24.031-24.183 24.302-24.405 24.573-24.673 25.191-25.235 25.281-25.579 25.850-25.982 26.136-26.247 26.284-26.784 27.214-28.100 28.400-28.625 28.690-28.854 29.514-29.779 29.886-30.228 29.576-31.103 33.313-33.812 34.922-35.200 35.586-35.760 35.909-36.518 36.771-37.543 37.660-37.877 37.941-38.112 38.389-39.382 39.451-39.892 40.898-41.135 42.064-43.245 45.731-47.511 48.778-49.540 50.734-50.921 50.034-51.802 52.478-52.800 52.897-52.947 53.057-53.527 56.113-56.584 57.691-58.027 60.850-61.188 62.359-63.479 63.821-64.613

| 22

7 Late Cretaceous-Cenozoic GPTS

Odin et al. (1991). (5) An age of 46.8 +__ 0.5 Ma by Bryan and Duncan (1983) correlated to C21n by Berggren et al. (1983a). (6) An age of 55 Ma for the NP9/NP10 nannofossil boundary (Swisher and Knox, 1991) correlated to the Paleocene/Eocene boundary (Berggren et al., 1985b). (7) Cretaceous-Paleocene boundary at 66 Ma (Harland et al., 1990). (8) Campanian-Maastrichtian boundary at 74.5 Ma (Obradovich and Cobban, 1975; Obradovich et al., 1986) based on the correlation of this stage boundary to the late part of C33n (Alvarez et al., 1977). (9) CampanianSantonian boundary at 84 Ma (Obradovich et al., 1986; Alvarez et al., 1977). In the revised version of their time scale, Cande and Kent (1995) adopted the astrochronological estimates (Shackleton et al., 1990; Hilgen, 1991a,b) for all Plio-Pleistocene reversals, and the 65 Ma estimate (as opposed to 66 Ma) for the Cretaceous-Tertiary boundary (Swisher et al., 1992). Otherwise the absolute age tie points are as for Cande and Kent (1992a) (Table 7.1, Fig. 7.6). Wei (1995) has taken issue with calibration points (2), (3), (4), (5), and (6) used by Cande and Kent (1992a, 1995). In place of calibration point (2) for C5Bn, Wei (1995) proposed a calibration point at 9.67 Ma for C5n (Baksi et al., 1993a) and 16.32 Ma for C5Br (Baksi, 1993). In place of calibration point (3) for the middle part of C6Cn at the Oligocene-Miocene boundary, Wei (1995) proposed a calibration point at 28.1 Ma for C9r (Odin et al., 1991). In place of calibration point (4) for the EoceneOligocene boundary, Wei (1995) proposed a calibration point of 35.2 Ma for C15r. This is based on a linear regression of nine age determinations in the C13r to C16n interval. In place of calibration points (5) and (6), Wei (1995) proposed a calibration point of 52.8 Ma for C24n.lr. This is based age within a polarity chron interpreted as C24n.lr in on an 4~ terrestrial sediments in Wyoming (Wing et aL, 1991; Tauxe et al., 1994; Clyde et al., 1994). The resulting time scale differs significantly from the Cande and Kent (1992a, 1995) time scales, particularly for the Miocene (Table 7.1). Obradovich (1993) has reviewed 4~ age control on Late Cretaceous ammonite zones in the U.S. Western Interior. These zones can be correlated to the European ammonite zones to provide the best available numerical age estimates of Albian/Cenomanian and younger Late Cretaceous stage boundaries (Table 7.2). Building on earlier work (De Boer, 1982; Schwarzacher and Fischer, 1982; Weissert et aL, 1985; Herbert and

Figure 7.6 Geomagneticpolarity time scale for Late Cretaceous and Cenozoic (after Cande and Kent, 1995). Polarity chrons with duration less than 30 ky are omitted.

] 23

7.2 Numerical Age Control

Pleistocene Pliocene

1

-0 m

Ma

2 2A 3 3A

4A 5

cC~ 0

Miocene

Z

--10

5A 5B 5C 5D 5E 6 6A

--

20

--

30

--

40

--

50

--

60

--

70

6B 6C

Oligocene

9 1011.12 13 15161718 19

0

20-

Eocene

13.

21 2223242526-

Paleocene

2728293031

32-

33-

Late

-- 80

34-

-- 90

Table 7.2 kte Cretaceous Stage Boundaries

Stage boundary

Polarity chron boundary

CretaceousRertiary Base C32r.lr Campanian/Maastrichtian Base C33n Base C33r SantonianKampanian ConiaciadSan tonian TuronianKoniacian CenomanianRuronian AlbianXenomanian AptiadAlbian BarremiadAptian

Cande and Kent (1992a) (Ma)

Obradovich (1993) (Ma)

66.00 73.78 74.50 78.78 83.00 84.00

65.4 t 0.1

Herbert et al. (1995) (Ma)

65.00 73.62 74.50 79.08 83.00 84.00

71.3 f 0.5 83.5 f 0.5 86.3 +- 0.5 88.7 t 0.5 93.3 f 0.2 98.5 f 0.5 112 f 1 121

For Cande and Kent (1992a) and Cande and Kent (1995). italics indicate calibration data for spline fit.

Cande and Kent (1995) (Ma)

99.5 111.6 122.0

7.2 Numerical Age Control

| 25

Fischer, 1986), Herbert et al. (1995) have determined the durations of the Aptian, Albian, and Cenomanian stages on the basis of cyclostratigraphy in the pelagic sediments in central Italy. A consistent bundling of limestonemarl couplets in a nearly 5:1 ratio, observed in Barremian-Cenomanian Italian sections, first led De Boer (1982) and Schwarzacher and Fischer (1982) to propose that the carbonate rhythms represent precessional forcing grouped into 100 ky increments by the eccentricity envelope. Herbert et aL (1995) combined cyclostratigraphic estimates for the duration of the Aptian, Albian, and Cenomanian with Obradovich's (1993) estimate for the Cenomanian-Turonian boundary (93.5 _+ 0.2 Ma) to calculate ages for the Barremian-Aptian, Aptian-Albian, and Albian-Cenomanian boundaries (Table 7.2). One clear discrepancy concerns the Campanian/Maastrictian boundary (see Table 7.2). This stage boundary is generally considered to coincide with the top of the G l o b o t r u n c a n i t a calcarata foraminiferal zone in Italian pelagic limestone sections, and this zonal boundary lies in the upper part of C33n. Therefore, the Cande and Kent (1995) time scale implies an age >73.62 Ma (age of top of C33n) for the Campanian/Maastrichtian boundary, whereas the definition of the stage boundary in the Western Interior (at the base of the Baculites eliasi zone) yields an age of 71.3 Ma (Kennedy et al., 1992; Obradovich, 1993). The younger age for this stage boundary appears to be consistent with Campanian foraminifera-bearing limestones from Mexico (Renn6 et al., 1991). In the deep-sea fan deposits of the Point Loma Formation (San Diego, California), the Campanian/Maastrictian boundary lies in C33n and can be defined on the basis of ammonites and molluscs (Bannon et al., 1989). Due to the endemic nature of the macrofauna, direct correlation from the Point Loma fauna to the Western Interior fauna is not possible.

Paleogene and Miocene Marine Magnetic Stratigraphy

8.1 Miocene Magnetic Stratigraphy Research on the magnetic stratigraphy of the Miocene began with the study of conventional piston cores with relatively low rates of sedimentation (Hays and Opdyke, 1967). In central Pacific cores, Foster and Opdyke (1970) obtained magnetic stratigraphies for sediments as old as Chron C5N (previously known as Chron 11) (Fig. 8.1). These studies pushed the limits of depth, and hence age of sediment, that could be reached by conventional piston cores. In the absence of suitable sections exposed on the continents, other methods of extending the polarity sequence and correlating it to the micropaleontological record had to be found. Consequently, studies were initiated on piston cores from the central Pacific in which hiatuses were present. These studies resulted in the correlation of siliceous fossil zones to sediments as old as Miocene (Opdyke et al., 1974; Theyer and Hammond, 1974). The magnetic stratigraphy of the Miocene languished until the development of the hydraulic piston corer (HPC) and advanced piston corer (APC) by the Deep Sea Drilling Project (DSDP) and the Ocean Drilling Program (ODP). The magnetostratigraphic coverage of the Late Miocene is relatively complete and high quality data are available from land sections (Langereis et al., 1984; Krijgsman et al., 1994a; Hilgen et al., 1995) and HPC/APC cores from North and South Atlantic (Fig. 8.2, columns 1-2-3-4-5), Indian Ocean (column 7), and Pacific Ocean (columns 8-9). As a result, correlations for this time period are available for both high- and low-latitude sites. The magnetobiochronology for the late Miocene was established by Berggren et al. (1985b) and revised by Backman et al. (1990) and Berggren et al. (1995). 126

8.1 Miocene Magnetic Stratigraphy

DECLINATION

Depth (m)

01 Brunh~ 1 (1) 2

~

3.

Matuyama ..... (_2_)_

0 1

(~ 180 I

~

J

Gilbert (4)

n

3Ar --

3An (5) 11-

| 27

~:

"~

(s)

. . . . . . . . . . . . . .

3B (7)

,

4n 4r (e)

4An (9) . . . . . . . . . .

4Ar

(lo)

. . . . . . . . .

~

....

~-~

5n (11)

Figure 8.1 Magnetic declination in RC12-65 after alternating field demagnetization at peak fields of 5 mT. The black (normal) and white (reverse) bar diagram on the left indicates the proposed extension of the geomagnetic time scale based on this early study (after Foster and Opdyke, 1970). Old polarity chron nomenclature given in brackets.

{')

Z Z_

(-) -I-

Z

"

I-

I::3 0

~ "z' - r ~rerrl--

z,,,

.

128

\

l

~Ill IlUl I II ! I \

'

I

I

I

I I

'

~

I

I I I

I

-..

m I

~

I

~

I

I

I

Ill

co

~

~

< ~ I- _

/

I I I

~

/

l

/

~ ~

I

I

~'

/

uJ ;

/ ~

,

=: j0, , -

oII11D~~II~I]

~

ill

[

Q

~.

(,.) -I u.I

~

I I I

'

I

I I

I I

'

I

I

:III fill

~

i.u

/

/

/

"

I

< > ~'

Z -'

i.u

l [ l / / l l m ~ EI l l t f l l l

Q

Ill!

Figure 8.2 Magnetic stratigraphy of Miocene sediments. For key see Table 8.1.

Table 8.1 Miocene Rock unit

l o Age H I

Region

A

@

NSF

NSI

NSA

M

1

Slte 608

Br-C6C

N. Atlantic

+42.84

-23.09

6

1.5

I

500

2

PotamidaKawllo SltC S63 Site 563 Site 643 Site 519 Hole 521A Site 704

GilLC4r

Crete

+35.12

+2S.20

7

325

3

75

Br-C9 Br-C9 Br-CSE Br-CSn

N. Atlantic N. Atlantic N . Atlantic S . Atlantic S. Atlantic S. Atlantic

+33.64

41.77 43.77 +.(HII -11.66 - 10 26 - 47 I

I 1 1 1 1 1

5 m .5 m .S-.S rn .I-.4 m .07-.1 rn 1-.05 rn

I I I 1

2W 200 270 303 37 430

I 2

.2S m

1

P

3 4 S

6 7

4A-5C

Br-CSAN

+33.64 +67.28 -26.14 -26.07 +007 9

I I

D

RM

DM

AD

A

ZB

LA

ZB

NCh

%R

I

RY

17

43

D-I

26

3

46 46 I45 34 20 4x

IX 1X I6 9 9

IS

41 41 24 34 5 4 50

8

-

4

-

-

2

-

-

4

R+

5 5

-

-

-

-

-

-

4 4

-

6

-

A A

Z-V ZB

I I

47 54

22 I6

45 29

V K

A

ZB

I

42

13

62

-

A

F%

D

2R

10

SX

-

71 41

IS

43 52

48 16 i

9 7 2

II

3

I I

VDI I

0

-

-

-

FCT

4

- I - I

-

-

I I

4

RT

B B A Z-V A-T Z B A-T Z B A Z B

-

A A

NMZ

A

-

-

5

References Clement a n d Robinson (1986) Langereis PI a/. (19x4) Khan er a/. (198.5) Miller rr a/. (19x5) Bled (19x9) Tauxe C I a / . (19x4) Heller rr a/. (19x4) Hailwood a n d Clement (1991)

Hole 6 9 6 8 9 Site 710

-043

Br-C6AA

S. Atlantic Indian

-.OM

-6l.X +59.1

Br-SAN

Pacific

-26.1

+161.22

2

1.2 m

1

310

s i t e 575

Br-SD

E. Pacific

+5.85

-135.04

4

.S m

1

130

12 Site 57X 13 RC12-65

Br-SB Br-Sn

Pacific Pacific

+33.93 +4.65

+ISM3

.I-.2 m

I

-144Y7

I I

.I m

I

160 24.2

Gu-CSAB CSR-CSB C5n-CSAr C3ANC4.2N

E. Pacific USA

+9.5x -Y4.59 +35.17 -120.25 +IX.75-77.90 +35.12 +2S.20

2 1 3 4

P

8

10 Site SUX

M-C17

~

150 124

Speiss (1990) Schneider a n d Kent (1990h) Barton a n d Blormendal (19x6)

II

5

Weinrich a n d Theycr

-

-

1

-

-

3

Heath er a/. (198.5) Foster a n d O p d y k e

(1985) 11

( 1970)

Site 845 Monterey Fm Buff Bay Crete

Cretc

X9 6X

.3 m

12 2 1

160

300 130 ho

-

F~K -

Ar

-

A

-

T

-

A

K

T

2-6 2-P 2-P Z

D-I F-V -

D-I

4

9 49 R + 5 7 -

6

l

-

-

F+ -

-

5 X

5 5

Schneider (1995) Omar7i e r a / . (1993) Miller er a/. (1994) Krijgsman el a/. (1994a)

] ~0

8 Paleogene and Miocene Marine Magnetic Stratigraphy

Unquestionably, the best results from calcareous sediments of middle and early Miocene age were obtained by Clement and Robinson (1986) from the North Atlantic at DSDP Site 609. The magnetic stratigraphy from this site is the most complete available for the Miocene. As a result, ancillary studies such as 8180 and 87Sr/86Sr stratigraphy (Miller et al., 1991b) have been carried out. Unfortunately for Miocene magnetobiochronology, the microfauna and flora at this site are not representative of the tropical Miocene from which the standard foraminiferal and phytoplankton zonations are derived. Miller et al. (1991b) have, however, detected rare tropical foraminifers in this core by processing large volumes of sediment, enabling them to correlate the fauna to the low-latitude zonation. The Miocene magnetobiochronology of Berggren et al. (1985b) relied heavily on South Atlantic cores from DSDP Leg 72 (Berggren et al., 1983b). Unfortunately, magnetostratigraphic data were never published; only the interpretation was published as black and white bars. It is therefore impossible to assess the quality of the magnetostratigraphic data, which makes the magnetobiochronology suspect. The Indian Ocean magnetostratigraphic record recovered at ODP Site 710 is of good quality, although hiatuses are present in the record (Schneider and Kent, 1990b). Correlation of low-latitude foraminiferal and nannoplankton zonations to the GPTS at this site has led to a revision of the Berggren et al. (1985a) magnetobiochronology (Backman et aL, 1990). These correlations have been confirmed by a recent study of Miocene sediment from the equatorial Pacific at ODP Holes 845A and 845B (Schneider, 1995). The Pacific magnetic stratigraphies were obtained using the ship-board pass-through cryogenic magnetometer at a 10-cm sampling interval after alternating field demagnetization at a peak field of 10 mT (Fig. 8.3). The correlation of the Miocene biozones to the GPTS is illustrated in Figure 8.8. The Oligocene-Miocene boundary and its position with respect to the GPTS are matters of conjecture. This is due to the uncertainty in correlation of the type sections for the Aquitanian and Chattian stages, where an unconformity appears to coincide with the boundary. The problem was discussed extensively by Berggren et al. (1985b). The boundary was placed by them within C6Cn. The bioevent often used to denote the boundary is the first appearance of G. kugleri. Lowrie (1989) pointed out that the first appearance of this species appears to be diachronous between the Mediterranean and the South Atlantic. Backman et al. (1990) have suggested that the Oligocene-Miocene boundary be placed in Chron C6Cr; however, this placement is by extrapolation and not by direct correlation to the GPTS. Cande and Kent (1992a) have placed this boundary at the base of C6Cn.2 and this determination is followed here.

8.1 Miocene Magnetic Stratigraphy

mcd (meters)

GPTS

Hole 845A DEC INC

160

13 ]

I1'r ~

Hole 845B DEC

INC

INT

~L~~ ---~ ~ ~ .

165

170

175 ~

180

185

190

195

200

_

2O5~ 210:

2151

> :?-'m

I > ;-_--~_

220 ~_

2251

23C

235

Figure 8.3 Magnetostratigraphic results from ODP Site 845 (after Schneider, 1995).

132

8 Paleogene and Miocene Marine Magnetic Stratigraphy

8.2 PaleogeneMagnetic Stratigraphy Unlike the record from the Miocene, which has depended heavily on marine cores, the correlation of the magnetic stratigraphy to the biostratigraphy of the marine Paleogene was first accomplished in a series of sections in north central Italy by Alvarez et al. (1977), Lowrie et al. (1982), and Napoleone et al. (1983), in a series of papers that have become classics. The pelagic limestones in these sections at Gubbio and Contessa have magnetization carried by magnetite and, in some cases, hematite. Both these magnetization components appear to have been acquired early in the history of the sediments. The antiquity of the magnetization has been verified by positive fold tests. Biostratigraphy in these sections required the use of new analytical techniques. Unlike most studies on deep-sea cores, the foraminifera in the indurated sediments at Gubbio had to be identified mainly in thin section (Premoli-Silva, 1977). Initially the nannofossils were considered too poorly preserved to be adequately identified (Premoli-Silva, 1977). It was not until the work of Monechi and Thierstein (1985) that a nannofossil biostratigraphy was available in these sections, although the poor nannofossil preservation in this area remains a hindrance to the precise correlation of nannofossil events to the GPTS. In the Contessa section, volcanic air fall ashes have been dated by the 4~ method (Montanari et al., 1988). The magnetic stratigraphy from the Contessa quarry section is shown in Figure 8.4. The sections of the region are well exposed in road cuts and could serve as type sections for integrated magnetobiochronology of Paleogene through Miocene time. The highest frequency of reversals of the geomagnetic field in the last 100 My occurs in the Miocene, inhibiting correlation of Miocene magnetic stratigraphies to the GPTS. The reversal rate decreases going back in time from the Late Oligocene to the Paleocene, facilitating the correlation of magnetic stratigraphies to the GPTS. The better quality magnetic stratigraphies for the Paleogene are plotted in Figure 8.5. The polarity zones from the Gubbio section have been given letter designations such as N+, which correlates to Chron 26n. In the time interval covered by the Italian sections, 33 normal polarity intervals occur in the GPTS, whereas 28 normal polarity zones are observed in the Italian sections. The discrepancy of six normal intervals occurs in the Chron 15 to 18 interval. Further sampling in the Upper Eocene may reveal the fine structure of this part of the magnetostratigraphic sequence. The magnetic stratigraphies from South Atlantic DSDP Sites 522, 523, and 524 (Tauxe et al., 1983c) cover most of the Paleogene except for a missing interval in the Eocene from Chron 21 to 23, representing approximately 6 My. The quality of the magnetic stratigraphy in these cores is

8.2 Paleogene Magnetic Stratigraphy

| 33

BIOSTRATIGRAPHY

Planktic Foraminifera

L! THOLOGY

Zones Gr. cerroazulensls

--X~._

--

--~" ~"

150~

,

--~ -- ~

Events

CONTESSA H I G H W A Y S E C T I O N

AGE POLARITY

G/o blgerina theka sem.nvoluta

Trunc orot al oides i

~,~N~

L A

rohm

O. beckrnanni

_1 -

o

VGP LATITUDE 0

90N

......==:==

Globlgemnatheka semtlnvoluta Orbulinoldes beckmanm

~'

90S

E M

~.

Globorotaha

~ '

_

0 i

lehnert

D

K

-

~n '

subconglobata

"~ ~"

aragonensls

C D

99m

_

-

~

-

Z':~

~

50-- ~

CJ3

aragonensls Hantkenlna

o,oo . . . . .

i~ ~!= . ~ : c~ :

~

I G. edgari .-,.,-G. velascoensls ~ ~ G, pseudomenardii

--r-

N

"

E

E A

G/oborotaha aragonensis to Globorotalta subbottnae

;

~'~

_

..

pentacamerata

--r-" ~um; --~- ~ -

E E

Hantkenlna 82

~"

el G/oborotaha

L

~

.,., O

Globlgertnatheka

R L Y

Globorotalia

_]

edgari

_

~

-

- G.-cmcin~a . . . . . . ! Globorotalia ~ - - G. t r i n i d a d e n s l s - - - ~ pseudobul/oides u~ ~ G pseudobulloides~

G. pusi//a pusH/a 0

O--

--

-~-

- .__.===~ ,.~,-------~

._

GI. eugubina

-

E

Globtgerlntds

Figure 8 . 4 Magnetic stratigraphy and biostratigraphy of the Eocene part of the Contessa

section (after Lowrie et aL, 1982).

exceptional, and their correlation to the Gubbio/Contessa sections and to the GPTS is unambiguous (Fig. 8.5). Farther to the south in Antarctic waters on the Maud Rise (ODP Leg 113), excellent core recovery at Holes 689 and 690 has facilitated the acquisition of high-quality magnetic stratigraphies (Speiss, 1990). Hiatuses hinder correlation of the magnetic stratigraphy to the GPTS in some intervals; however, excellent results were obtained for Late Oligocene to Middle Eocene time (Chron 8 to Chron 21 in Hole 689B). An excellent record was also obtained in Hole 690B from Late Paleocene (Chron 26) to Early Eocene time (Chron 24). A short magnetostratigraphic section has been reported from Hole 702B by Clement and Hailwood (1991) in sediments of Middle Eocene age (Chron 18-21). Other data are available but are often discontinuous (Fig. 8.5). It is interesting to note that almost all of these data come from the Atlantic Ocean and Mediterranean area. Data are lacking from the Pacific and Indian oceans. Data are available from the middle Eocene of the Atlantic margin (Miller et al., 1990), southern England (Townsend and Hailwood, 1985), and from

Ma

SOUTH

ITALY

_ _ . , . ,,.. OL

25

1

..,,

ATLANTIC

8

L A IT

m

GE

m

O--

m

C

30'

E

m

A

m

ER 12

N L E

Y

..-m

....o

.~

L A T E 17 n

I

E-

o

40'

M I

.,, . . . . - ,-,

Co

14 Z§

D E

45"

y-

L

E

X§

N

I E E

;

A

'

n 23_. . . . .

W-

V+

R UL

'~ 7

Y

551:24 r R§

pL k 25 r

A A L T

i

i, 'k

E E 601:26 r

C

Q-

1

i

Ciz 27 r 28 n

E!A

N§

IR

NiL ~ ~ ElY

65

L

b

% H§

Figure 8.5 Paleogene marine magnetostratigraphic records. For key see Table 8.2. 134

Table 8.2 Paleogene Rock unit

Lo Age HI

Region

NSE

h

NSI

NSA

M

D

RM

AD

A

NMZ

NCh

%R

ILK I

T-A T

2-V 1-V

F-V D-I

31

14

51

I6

10

60

A T A A A A A

v

I 2 - V FV Z-B I 2-B I Z-B I 2-B I 2-B I

36 33 27

12

43 51 58 46 46

-

-

R+ R+ R+ R+ R+ -

F+ -

A A A T

2-B

I I 2-v I B V

21 31 3x 27

9 5 7 1 4 4 6 13 53 23 6X -

A T

V

-

7 33

13

DM

RT

FCT

Q

R+

F+

-

-

X 6

1 Scaglia 2 Scaglia

C6-C20 C17-C29

Italy Italy

+43.38 +43.37

+I256 +12.58

1 1

.3-6m 202

I I

140 A 150 -

3 Site 558 4 Scaglia 5 Site 523 6 Site 524 7 Site 524 X Site 522 9 Hole 702B

CSAD-C7 Clh-C29 Cll-C20 C23-C31 C23-C31 C6-CI6 CI 8-C2 1

N Atlantic Italy S. Atlantic S. Atlantic S. Atlantic S. Atlantic S. Atlantic

+37.77 +43.38 -28.55 -15.01 -15.01

+37.34 +12.56

2 3 I 1 I I 1

25-.4 m 3-.6 m .1-.4 m 1.4m 1.4m I-.4m PTM

1

30 - 160 A I 130 - 2 ! X 290 - -

10 Site 577 II Site 689 12 Site 690 13 Lincoln Creek

C23-C30 C7-CI5 CY-C26 C6C-CI7

N W Pacific S. Atlantic S. Atlantic USA

+32.44 +MI4 +003 +47.00

+ 157.72

2 2 2

.1-.5m

25 m

1 1

1

25 m Y7

1 3

14 Atlantic Coastal PI. I5 Alpine Scaglia

C8-C23 C24-C34

USA Italy

+39.44 +46.00

-74.72 +11.75

1 5

62 359

1 1

-

-

-

A-T

V

-

28

5 7 1 -

76.2

0

I

A

V

I

14

6 6 2 -

References Lowrie er a/. (1982) Napoleone rr a/. (1983)

I6

London Basin

CZI-C25

Alabama Coastal PI.

Cll-C16

England USA

-26.11 -50.95

+51

-31.50

-2.25 -3.51 -3.51 -.I3 -26.37

-M

-65 -123.5

I 1 1 1 1 1

ooo

10

152

2

-88.00

3

153

I

155 10

-

-

-

I

6 9 - 150 144 -

2900K 100s 127 -

-

-

2-V

ZB

I V

21 21 63 Y

13 9 9

Y 13

5

4

54 5 7

5

0 28

-

5 9

-

5 5 5 5

-

6

-

6 7 7 4

-

4

-

5

-

-

~

-

2 6

Khan el a/ (19x5) Lowrie el a/. (1982) Tauxe rr a/. (1984) Tauxe rr a/. (1984) Tauxe el a/. (1984) Tauxe el o/. (1984) Clement and Hailwood (1991) Bleil (1989) Speiss (1990) Speiss (1990) Prothero and Armentrout (1985) Miller el a/. (1990) Channel1 and Mediva (1981) Townsend and Hailwood (1985) Miller e r a / . (1993)

136

8 Paleogene and Miocene Marine Magnetic Stratigraphy

the late Eocene and Oligocene of coastal Oregon (Prothero and Armentrout, 1985). The magnetobiochronology of the Paleocene has been discussed by Berggren et al. (1985a), Aubry et al. (1988), and Berggren et al. (1995). The Eocene/Oligocene boundary has been placed by the International Stratigraphic Commission in the Massignano section (Italy) at the level of the last appearance of the foraminifera H a n k e n i n a (Montanari et al., 1988). This position would place the stage boundary at C13r(0.14), close to the disappearance level of H a n k e n i n a at DSDP Site 522 (Poore et al., 1984), but above this bioevent in the Contessa Quarry section. From the original paleomagnetic investigation at Massignano, Bice and Montanari (1988) postulated the presence of a short normal polarity subchron immediately prior to the Eocene/Oligocene boundary. This subchron was not identified by Lowrie and Lanci (1994) and the present consensus is that the position of the boundary should be placed at C13r(0.14), as advocated by Cande and Kent (1992a). The correlation of the Paleocene-Eocene boundary to C24r(0.6) (Cande and Kent, 1992a) is the same as that adopted by Berggren et al. (1985a) and close to that observed in the Umbrian sections by Lowrie et al. (1982). Berggren et al. (1995) give an uncertainty of about 1 My in the correlation of this stage boundary. The early/middle Eocene boundary is correlated to the GPTS in the Contessa sections and correlates to the top of C22n. The evidence for the placement of the middle/late Eocene boundary in the polarity sequence is not so clear, and this boundary is placed in late C18n in the Contessa Highway section. Berggren et al. (1995) choose to place this boundary in C17n on the basis of fossils from DSDP Site 523 in the South Atlantic (Poore et al., 1984). This placement is followed here; however, the correlation of this boundary to the GPTS depends on the paleontological criteria chosen to define it.

8.3 Integration of Chemostratigraphyand Magnetic Stratigraphy In Plio-Pleistocene sediments, 6180 stratigraphy provides a high-resolution stratigraphic correlation tool and a means of tuning the time scale and assigning absolute ages to polarity chrons (Ch. 6). The Cenozoic 6180 record derived from Atlantic benthic foraminifera (Fig. 8.6) (Miller et al., 1987) may be interpreted as changes in the temperature of ocean bottom water or ice volume changes, or both. Shackleton and Opdyke (1973) have shown that ice volume changes dominated the Pleistocene benthic 6a80 record. Separating temperature change from ice volume for the early Cenozoic record is more

| 37

8.3 Integration of Chemostratigraphy and Magnetic Stratigraphy

5 TM O (Atlantic sites) (0/00) 4

3

2

1

0

-1

0 Ii~_

I

"

Plio-Pleistocene

It II

10 §

I§

20

t~

30

O exl) ._.~-

9

'~

40

O Ud

50 benthic foraminifera 60

|

N

@

@o

Cretaceous

70 ice sheets i possibly ice free 0

2

I

I

4 I

6 I

"ice free" 8w = -1.2 %0 2

!

I,,

8 I

4

I

"modem" 8., = -0.28 %0 I

6

I

8

I

l0

I

12

I

14

Figure 8.6 Composite benthic foraminiferal oxygen isotope record for Atlantic DSDP sites corrected to Cibicidoides and reported to Pee Dee Belemnite (PDB) standard. Chronostratigraphic subdivisions are drawn after Berggren et al. (1985a). The smoothed curve is obtained by linearly interpolating between data at 0.1-My intervals and smoothing with a 27-point Gaussian convolution filter, removing frequencies higher than 1.35/My. The vertical line is drawn through 1.8%o; values greater than this provide evidence for existence of significant ice sheets. The temperature scale is computed using the paleotemperature equation, assuming Cibicidoides are depleted relative to equilibrium by 0.64%0. The lower temperature scale assumes no significant ice sheets, and therefore 0w = 1.2%o; the upper scale assumes ice volume equivalent to modern values, and therefore 0w = 0.28%0 (after Miller et al., 1988).

difficult. In the Paleocene, the 6180 values vary about a value of 0%o, then in the late Paleocene 6180 values begin to become more negative, peaking at -1%o in the early Eocene (Fig. 8.6). The 6180 values become steadily more

138

8 Paleogene and Miocene Marine Magnetic Stratigraphy

Atlantic Composite (0/00) b 18 Ma

~

, i

,

,

m

38:

-

4

15--

~ --

_

~

|

m

|

~5~AO

~ ,'---', E ,

I

I

1.5 I

1.0

'

I

.5

'

I

2.5

'

Mi3

SB

~

_

I

0

3A

. . . . .

' u.I ~

I

2.0

_1

,

5C

i

50

I

150~

(Mi2

2

1.5

N

6~

~,., Milb

-

20--

--~

~

.~

leA , r m ' '

i

25 - -_~ _

, tu m

-

z

-

~

'1n l

~', _~ ~ l

i

- 11

tu

z

| -

~Oi2a

I I

2

12 m - -

-

35-.,

I

==

--

6C

1 7 ' ~,-'. 8 9 10

I

~) , ..~ m

-

'

I

..

30--

6AA i 6B

I

i

13

115 I 16

~

0il " ~

Figure 8.7 Composite record from the Atlantic indicating position of oxygen isotope events for Oligocene and Miocene time (after Miller et aL, 1991a). Inset shows high-resolution record for Mi3-Mi4 (after Woodruff and Savin, 1991). Time scale from Cande and Kent (1995).

positive throughout the Eocene, increasing sharply at the Eocene/Oligocene boundary to values of about + 2.8%0. The ~180 values then oscillate throughout the Oligocene and early Miocene, and a sharp change to more positive values occurs in the interval from 15 to 13 Ma, and the values then oscillate around +2.25%o for the rest of the Miocene. These changing values of ~ 8 0 in benthic foraminifera have been interpreted by Miller et al. (1987, 1991a) in terms of bottom water temperature and ice volume changes. Miller et al. (1987) interpreted the changes ob-

~

~

'-.,o

-x_~

Mi1

I

0

Z

o3 Z (2

Z 0

CN3

CN2

I

CP

CN1

19

Z

z

Z

NN4

,

M3

N4

P

M2

, E

Aquitanian

g

.~

~

E

.i

rn

~ u~

~5

'

~

Lower ,

.r

I

(])'

E

0

._o

Miocene

I-'-

01 C6C

-

18

om I

C6A

i.~

17

' ~'

.

"-

,-

_,.o

I_ 9 .

ii

16

i, ~

oLU

~,o I

Q~

0

I.

C6

, o << - _ ~ <

i

m

~176

, VVgO I _r < I T "

..0

i-=

-(::~

I ,

I

-~

I '

I

A

F.

m-|

I

~ o

a.

, 8 II'~

I

,,~

, 0

0

<

O~

o

0 l-"

o

i5 I 12il\6

C5B lC5C C5D

Ch

I

9

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O_ C~

Middle

~oo
:z)

(1)

.I-m

O3

~ ~

.~.

> ,-~

Burdigalian

Lawhian

5

77

,

N5

~

~

M4

M7

9

N6

z

M5

N8

~

(x:) Z

,

e~

E

E ._~

=~

(,-

o 1=

~

N

r-~

,

lM

Serravalltan

Plto

NN2

N:

z

o

N7

N10

Upper

Stage

I

I

t.o Z Z

, z 2 , ~o~

c~

,-

~

y12

Tortonian

Zanclian Messinian

Mil a

,

o') Z Z

_

N16

M i3

' 14 MI

Mil b

Mi2

uo Z

I'.. Z (.)

,,-

~ "-

N17

Z

1

~

z

~

pL1

1995

,--

-

-

z t-

Berggren

Mi4 Mi3

z

z

~

z~ z,,-

"11

N191N18

m

SLUEJO-I

S

~

,

-

z~

z,-

& :

g

ct CN7 7; NN9

CN9

13 12

Mi7 Mi6 Mi5

I I I I I I

I

0

""

1971

Mi8

o

Mart'n'

o ~--

z

0

CNlO

1973-75

,-

souueN

Bukry

increasing b C values

Mi9

I

$ 2

o

"-.,o

I

Mil0

L

|

~

co

,.09

00

L_

0 0 0 0 0 0 0

0.7089 0.7088 0.70870.7086 0.7085 0.70840.7083-

19

21

20

23

22

24

' r r,~

1,~ r.b.O0c hi)

P,.,~

o N r.~

-~-~=

~,~

=

S r.~

~

~-

2~

N~

2~

NL

r~

q~

~o

I-

o ;~ o

o~d

,,,w,w

~ m

.,- " ~ ' ~

o

Figure 8.8 Polarity time scale for the Miocene (after Cande and Kent, 1995) correlated to foraminifera1 zones and nannofossil zones (after Berggren et al., 1995), to isotopic events recorded in 6IXOdenoted by the prefix "Mi" (Miller et al., 1991b). to 6I3C (after Woodruff and Savin, 1991). and "Sr/ %r (after Hodell and Woodruff, 1994).

140

8 Paleogene and Miocene Marine Magnetic Stratigraphy

served in Paleocene and Eocene sediments as largely due to changes in bottom water temperature. They argued that ice sheets became an important factor only after the Eocene/Oligocene boundary, and only since this time (---34 Ma) do oscillations observed in the 8180 denote ice volume change. This is supported in some cases by the covariance of 8180 values from benthic and planktic foraminifera. Times of sea level fall are presumed to be coincident with positive shifts on the post-Eocene 8180 curve, with changes to more negative values representing deglaciation. It should be noted that because of the smoothing employed in the 8180 record shown in Figure 8.6, it is not possible to see the high-frequency Milankovitch cycles characteristic of the Plio-Pleistocene record. Miller et al. (1991a) have presented a new and updated 8180 record for the Oligocene and Miocene (Fig. 8.7). They argue that continental glaciation began in Antarctica in the latest Eocene, at about 35 Ma. They have formally defined nine oxygen isotope zones in the Oligocene and Miocene which they have designated Oil and Oi2 and Mil to Mi7, ranging in age from Early Oligocene through early Late Miocene (Figs. 8.7 and 8.8). Miller et al. (1991a) interpret these isotope events in terms of glacial episodes. It is clear that this 8180 stratigraphy does not have the resolution of the Plio-Pleistocene record and isotopic events have age uncertainties of 0.5 My using the sampling strategies employed by Miller et al. (1991a). The question arises as to whether early Neogene and Paleogene records can, in the future, be used to yield a high-resolution 8180 stratigraphy. Woodruff and Savin (1991) have studied in detail the mid-Miocene 8180 shift at DSDP Site 574, and it is possible to correlate isotope stage Mi3 to the section at Site 574 in the 150 to 156 mbsf interval (Fig. 8.7). It would appear that the potential exists for an extension of high-resolution 8180 stratigraphy to sediments of Miocene age, and perhaps to the entire Cenozoic. Site 574 has no magnetic stratigraphy; however, the correlation to the Atlantic composite record is supported by the biostratigraphy. Shifts in the 813C record have been utilized in sediment correlation (Berger, 1982; Haq et al., 1980; Miller et al., 1988, 1989). Changes in 813C are thought to be due to burial of organic carbon, erosion of carbonate sediments (reservoir exchanges), variations in biomass (climate change), or changes in productivity (Berger et al., 1981). For the Pleistocene, inphase changes of 6180 and 613C, with 8180 becoming heavier and t~13C becoming lighter during glacial episodes, are well known (Shackleton, 1977). Such in-phase changes are common throughout the Neogene beginning by Middle Miocene time (Woodruff and Savin, 1991). It is interesting to note that for two earlier 8180 events thought to represent glacial events (Oil and Mil), the t~13C shift is synchronous with the t~180 shift, but toward heavier values, not lighter ones.

OiZa

012

011

Figure 8.9 Polarity time scale for the Paleogene (after Cande and Kent, 1995) correlated to foraminifera1 and nannofossil zones (after Berggren et al., 1995). oxygen isotope events indicated by arrows (after Miller et al., 1987. 1991a), and X7Sr/XhSr record modified after Mead and Hodell (1995).

]~

8 Paleogene and Miocene Marine Magnetic Stratigraphy

The best documented 613C carbon shift is the prominant upper Miocene 6~3C shift of -0.7%0 (Keigwin, 1979) which coincides with Chron 3Bn at about 7 Ma (Haq et al., 1980) (Fig. 8.8). This carbon shift seems to occur not only in the marine record but also in ancient soils in the land record (Quade et al., 1989). The land and marine occurrences are both correlated to the GPTS. Recent studies of ODP cores off Greenland indicate that this event may correlate with the onset of glaciation on Greenland (Leg 152 shipboard party, 1994). This event is, therefore, an important link between the marine and nonmarine records. A second important 6~3C shift of about + 1.5%o close to the boundary between the lower and middle Miocene has been documented by Vincent et al. (1985). The middle Miocene is marked by the most positive 6~3C values of the entire Miocene and this interval has been called the Monterey event (Vincent and Berger, 1985). This 6~3C shift is hypothesized to be the result of organic carbon being locked up in sediments around the Pacific Rim. Woodruff and Savin (1991) have studied the 6~3C fluctuations in the middle Miocene in detail and have designated 6~3C isotope stages I to IX. The relative positions of the boundaries of carbon isotope stages I through VI are shown in Figure 8.8. The cores which record these stages do not have magnetic stratigraphies and the biostratigraphic zonation is nonstandard, which makes correlation difficult. Nevertheless, future studies will undoubtedly result in the 6~3Crecord being a useful correlation technique for Miocene time. Miller et al. (1988) have provided a 6180 and 613C record for the Oligocene from DSDP Site 522, which has good magnetostratigraphic control. They suggest that this record serve as a type record for isotopic stages in the Paleogene. Two prominent 6~3C excursions are seen in this core which are correlative with the 6~80 events Oil and Mil (Fig. 8.7). One of the most promising isotopic correlation techniques to emerge in the last decade is the use of the changing ratio of 87Sr/86Sr with time in ocean water (DePaolo and Ingram, 1985). It is particularly useful from the beginning of the Oligocene to the present, when 87Sr/86Sr in seawater changed from .7077 to .7090 (Figs. 8.8 and 8.9). High-resolution data are available for Miocene and younger sediments (Hodell et aL, 1991). Miller et al. (1988, 1991b) have studied deep-sea cores with magnetic stratigraphy to create a record of Sr isotope changes for the Miocene that can be correlated to the GPTS. The study by Hodell et al. (1991) used a combination of magnetostratigraphically dated cores and data correlated to the GPTS indirectly through biochronology. For some time periods where there is rapid change in 87Sr/86Sr, such as the Lower Miocene, it is possible to correlate magnetostratigraphic sections to the GPTS using the 87Sr/ 86Sr record.

8.3 Integration of Chemostratigraphy and Magnetic Stratigraphy

143

The reason for the strontium isotopic ratio changing with time is a matter of debate. Strontium enters the ocean through the erosion of mountain belts, ancient shield areas, and oceanic ridge systems. Fortunately, it is not necessary to know the cause of the variation in order to use it for stratigraphic correlation. The integration of chemostratigraphy with biostratigraphy and magnetic stratigraphy can enhance the precision of stratigraphic correlation in the Cenozoic (Figs. 8.8 and 8.9). Berggren coined the term magnetobiochronology, which might now have to be amended to magnetobiochemochronology, which is a little awkward; but MBC chronology might be acceptable.

Cenozoic Terrestrial Magnetic Stratigraphy

9.1 Introduction One of the long-standing problems of stratigraphy has been the correlation of nonmarine sediments containing vertebrate fossils to marine sediments containing invertebrate fossils. Due to correlation difficulties, the stratigraphy of marine and nonmarine sediments has developed more or less independently. As geomagnetic polarity reversals are globally synchronous, magnetic polarity stratigraphy offers the opportunity to correlate between these two contrasting sedimentary environments. Magnetostratigraphic studies of terrestrial sequences began in the early 1970s in a Plio-Pleistocene mammal-bearing sequence of the San Pedro Valley, Arizona (N. M. Johnson et al., 1975). This study, subsequently updated by Lindsay et al. (1990), demonstrated the potential of magnetic stratigraphy in terrestrial sequences. The authors were able to correlate widely separated fossil localities and place them in stratigraphic sequence. Correlation of polarity zones to the GPTS provided the time frame. The sediments of the basin were mapped using polarity zones and the rates of basin subsidence determined. In the 20 years since the publication of this paper, magnetostratigraphic studies of nonmarine Cenozoic sediments have become commonplace and have been used to solve problems in vertebrate paleontology, faunal migration, sedimentology, basin subsidence, and tectonic history. Correlation of Plio-Pleistocene land mammal sequences to the GPTS has been accomplished in North America, Western Europe, southern Rus144

9.2 North American Neogene and Quaternary

145

sia, China, the Indian subcontinent, South America, and Africa (Fig. 9.1, Table 9.1). First and last appearances of mammal taxa have been correlated to the GPTS in many individual sections, some of which have radiometric age control. Magnetic stratigraphy of Plio-Pleistocene lake and loess sequences has been motivated by the desire to understand continental climatic response to Plio-Pleistocene glaciation. Good quality Plio-Pleistocene lake sediment magnetic stratigraphies exist for North America, Japan, Australia, and Europe (Table 9.2). The climatic record from lakes is often provided by downcore changes in pollen type and abundance. For loess, good quality magnetic stratigraphies have been obtained from Europe, Central Asia, North America, South America, and, most importantly, China (Table 9.3).

9.2 North American Neogene and Quaternary For North America, magnetic stratigraphy of mammal-bearing sequences covers almost the entire Cenozoic and extends into the Cretaceous. The terrestrial mammal-bearing stratigraphy in North America is divided into 19 land mammal ages (LMAs) based on stage of evolution and the appearance of immigrant mammal taxa from other continents, usually Eurasia, or at the end of the Cenozoic, from South America. Magnetic stratigraphy offers an independent means of ordering events in mammalian biochronology and thereby enhancing studies of tempo and mode in evolution. The Plio-Pleistocene nonmarine magnetostratigraphic record is well known for most continents and the correlation of the mammalian zones to the GPTS is reasonably well established (Fig. 9.1). In North America the record is more or less complete, except for the Late Pleistocene and the earliest Pliocene. The boundary between Blancan and Irvingtonian fauna occurs at the top of the Olduvai subchron, with the first appearance of Mammuthus in North America. The base of the Blancan LMA is tentatively placed in the Sidufjall subchron at ---4.8 Ma. Figure 9.2 shows the correlation of first and last appearances of mammal taxa to the GPTS in North America from Late Miocene time. The first appearances may be evolutionary first appearances, such as that of Equus at Hagermann in Idaho, or migrationary first appearances such as that of Mammuthus (elephant) just above the Olduvai subchron at Anza-Borrego in southern California. Considerable progress has been made in correlating the Miocene LMAs to the GPTS in North America (Fig. 9.3, Table 9.4). Initial studies on Miocene sediments by MacFadden (1977) and Barghorn (1981) have been followed by recent studies by Whistler and Burbank (1992) and MacFadden et al. (1990b) that extended this correlation to sediments containing Clarendonian and Barstovian fossils of middle and early Miocene age. Whistler

146

9 CenozoicTerrestrial Magnetic Stratigraphy

and Burbank (1992) placed the boundary between the Clarendonian and Hemphillian in early C4Ar and the preceeding boundary between the Barstovian and Clarendonian at the base of C5Br. MacFadden et al. (1990b) placed the Hemingfordian-Barstovian boundary within chron C5Br. The boundary between the Arikareen and Hemingfordian has been placed within C5Er by B. J. MacFadden (personal communication, 1995).

9.3 Eurasian Neogene Eurasia has relatively long Neogene terrestrial sequences with abundant fossil content. The most successful studies have been concentrated along the southern deformed margin of Eurasia where Cenozoic sediments have been exposed by Alpine-Himalayan orogeny. These areas extend from Pakistan and India, where much work has been done on the Siwalik sediments particularly in Pakistan, through the Caucasus and Turkey to the Mediterranean and Spain. Unfortunately, many of the areas of Western Europe where classical faunas were initially described are poorly suited to magnetostratigraphic study because of limited exposure and, in some cases, lack of knowledge of the precise location of fossil finds. The European early Pliocene land mammal age, the Ruscinian (Fig. 9.1), is younger than the flooding of the Mediterranean at the MiocenePliocene boundary, which is now well-dated magnetostratigraphically and astronomically (Ch. 6). Magnetostratigraphic studies in Spain (Opdyke et al., 1989, 1996) place the Turolian-Ruscinian boundary (=Mn 13/14 boundary) in the early Gilbert, at about the level of the lower boundary of the Sidufjall subchron. The Vallesian-Turolian boundary (Fig. 9.3) is placed at 9 Ma by Berggren et al. (1985b). This is close to the placement of the ClarendonianHemphillian at 8.8 Ma (Chron 4A) (Tedford et al., 1987; Whistler and Burbank, 1992). The position of the Astaracian-Vallesian boundary has been a subject of controversy because the base of the Vallesian is defined by the first appearance of Hipparion, several species of which migrated to Europe from North America. The appearance of Hipparion in the Siwaliks of Pakistan is well dated magnetostratigraphically at 10.8 Ma (Johnson et al., 1982). On the other hand, Berggren et al. (1985b) preferred an earlier age of 12.5 Ma based on radiometric ages from Germany that correlate with

Figure9.1 SelectedPliocene and Pleistocene terrestrial records fromNorth America,Europe, Pakistan, China, Africa, and South America. Numbersrefer to individual studies (see Table 9.1). Time scale from Cande and Kent (1995).

ILl

I II

-r

~

~'~'~

I

I

nun

O ~0

% %% % % %

I I <

I

Z

C)

,, /

B

i /

II

UJ (1. O IJJ %

ii

om

%% %

%

! %

o-

-i eo!JJV ueis!)led eu!qo e!ssn~l'S edoJn3" M

%/

% % %

ue!uoJleN

ue!uexJnl

uepeuesu:l

eo!Je'"V'S

f

%"~

% % %

["

ue!Jelv ue!nbFI

] ~

ue!lodseJ!s lue,Jele 0 -uu!eq u!e~s 9

eo!JeuuV "N

ue!ueMeqfN ueLue~

3N3::)01SI

~i L

3"Id

ueeqnoA I eul^Ol)l -!qeqeA)l ue,qoueJ-le . nn!A

l essePO I ,~Jdeq]4 ~" ,'j"

ue!uoIBu!AJl _o NL

ue!Jenb~en H

loJles

Jo!u!d ue!pnoxnoqz

ue!JeH

uesouJ Jeqe~,UOlN

~ .~ N~

]

~1~ u e e l 6 u ! r

~Iueepoe8

uetAeplolN J',l ue!u!osnu

~1~

ueoueI8

El ~

NV~:~

3N:iOOl'ld

ue!loJn.L ue!ll!qdmeH

I:1V i;

8

zc~'N

~"

I:Is

3N3DOI

Table 9.1 Plio-Pleistocene Mammals Rock unit

Lo Age Hi

Region

A

1 St. David Fm.

B-Gil

USA (A)

+32

2 Vallecito-Fish Crk Vallecito-Fish Crk

M-GiI M-GiI

USA (Cal) USA (Cal)

+33.15 -116.35 t33.15 -111.35

3 Jaw Face 4 Glen's Ferry

B-M M-GII

-50.66 - 107.87 +43 -115.29

M-Ga G-4n M-BiI B-Ga M-3A B-M M-4N B-GiI M-Ga B-M M Ga GI1 B-M M-GiI

Canada USA (Idaho) USA (Ar) USA (Ar) Spain Pakibtan Pakistan China China Kenya Ethiopia Bolivia Argentina Bolivia Argentina Bolivia

t32.75 +34.55 +40.62 +32.93 +33.00 +36 +37 +04 +5 -21.5 -21.32 -19.75 -38 -16.5

Gil-4A B-GiI B-Ga

Argentina Australia Argentina

-27.33 -66.9 -38 +I44 -38.33 -57.66

5 6 7 8 9 10 I1

12 13 14 15 I6 17 18

111 Ranchbeds Verde Fm Teruel Basin Upper Siwalik Upper Siwalik Louchuan loess Yushe Basin Koobe Fora Shungura Tarija Frn. Uquia lnchasi Beds Enscnadense LaPaz-Calvono

Q,

NSE

-110.2

-109.5 -112 +.001

+73.73 +73.5 +109.23 +I13 +36.3 +36 -64.75 -65.35 -65.5 -57.57 -68.0

1 9

4 9 3 3 8 I 25 9 25 4

NSI

NSA

D

RM

A

AD

K J-K

A

B

F

1

2

3

62

R+

1 1

2 2

3 3

56 56

R-

2 3

78 85

R+

50 46 50 50 49 52 58 46 50 47 85 82 46 47

-

400

150 150

4000 4000

- J-I F -

A T

B B

D F

II 500

F -K J-l

A A

Z-P B

F F D-l Z-V V V-F F F-V D-I V V F F.V. V F-V D-I F-DI

357 176 35 I64 78 113 132 .3 m 280 ~

1 3

3 3 3 3-5 3 1 3 ~

310

1

100 20

5

54

1

129 100

3

29 80

3 3

1

A-T B A B T 2-P 1100 A B 1850 A B 150 T 810 T 2-P 290 A J-I A-T %-P 770 A X A V T-A 2-P 650 K B 200 K 120 - I T-A Z-P 14.5 - -~ A B 2ho K I-H T Z-B 100

450

120

NMZ NCh %R RT F.C.T. Q

DM

149

3 3 I 3

4

M

K K V

F J - J-ILK - ~

~

2 5 2 3 2 16 9 38 7 20 12 1 7 4

4 4 3 11

2 5 3 3 6 3 4 4 2 2 2 1

2 3

R+

R+

RR+ R+ R+

~

R+ R+ -

R+ -

References

N. M. Johnson er a/. (1975) 5 Opdyke el a/. (1977) 7 N. M. Johnson er a/. (1983) 7 Barendregt era/. (1991) 7 Neville er al. (1979) 7

5 6 7 5

8 4 6 8 7 7 4 8 3

R+

Galusha er a/. (1984) Bressler and Butler (1978) Opdyke el a/. (1996) Keller el a/. (1977) Opdyke era/. (1979) Heller and Liu (1984) Tedford er a/. (1991) liillhouse el a/. (1986) Brown el 01. (1978) MacFadden el a/. (1983) Marshall er a/. (1982) MacFadden er a/. (1983) Nabel and Valencio (1981) lhouveny and Servant (1985)

19 Corral Quemado Victoria Buenos Aires

2 12 2

.2rn

1

Yo0 24 26

K J-I K ~

-

T-A Z-B F-V A-T PCA V A-T B F-V

8 11

17

I 4 3

57

R+ -

56

~

9 6 4

Butler er a/. (1984) Whitelaw (1991a) Orguira (1990)

Table 9.2 Lake Sediments Rock unit Lake George Tule Lake Bonneville Basin Lake Tyrell Lake Biwa Searles Lake

Lo Age HI B-Ga 6-Ga B-Ga B-Ga BB-Ga

Region

h

a

Aust. (NSW) USA (CA) USA (Utah) Aust (Vic)

-35 08 +4L0 +4O.X 35.34 +35.25

+ 14Y.42

I

XY

-121.5

3 1

3x4 144

Japan

+35.72

-112.54 +142.XI + 136.27 - 117.34

NSE

3 I I

NSI

112 -XOIX)

120

NSA

3 I 1 I I 3

IM

D

RM

3 6 - - K K-A Ill K 22 - 200 K 700 - 370

DM

AD

A A

V B

-

-

A-T

V-7.

A

B B

A-I

A

I I

F D~I D-I I

NMZ

NCh

Y 17 3 X

3 3 2 4

3 5 41 7 47 -

I 23

1 3

i ) 47 --

%R

RT

1.C.T.

0

-

4

Singh u r d . (19x1)

5

Rieck CI u/. ( l Y Y 2 ) Eardlcy PI u / (lY73) An el a / (IYXh) Yaskawa er u/. (1Y73) Liddicoal rr d.( I Y X O )

I

-

5

-

4

-

4

Reference5

Table 9.3 Loess Records ~~~

Rock unit

Age

Region

A

Baoji Lanzhou Fairbanks Xifeng Luochuan Tashkent Balneario Lantian Shan XianXifeng Baoji Tajik

Ga-Gil B-M B-Ga 8-Ga B-M B-M B-Ga B-GiI B-M

Shaanxi-China Gansu-China Alaska-USA Gansu-China Shaanxi-China Tadzhikistan Argentina China Henan Gansu-China Shaanxi-China Tajikistan

+34.33 +35.2 +64.51 +35.7 +35.8 +41.20 -38.18 +34.2 +34.3 +35.70 +34.33 +38.33

B-Ga B-M

+107.01 +103.16 - 147.45 +107.60 + 109.2 +69.18 -57.64 +109.2 +111.9 +107.8 f107.1 +68.48

N.S.E.

NSI

1 1 6

576 40 283 1848

1

2 6 1 1 2

488 1500 180 500

1 7

516 250

NSA

226 1 1

M

RM

DM

AD

A

NMZ

190 330 32 200 150

K

T T A

90

K

20 195 135

K

T T T A T T

B B B 2-B Z-V B Z-V Z-V Z-B

D-I V I D-I D-I-F D-l D-l D-I V

12 4 10 21 7 3 7 21 6

186 100

K

T T-A

Z-B Z-P

D-I D-I

16 2

---

D

-

K

K-l

~

K

-

NCh

%R

RT

F.C.T.

Q

Evans er a/. (1991) Burbank and Li (1985) Westgate er a/. (1990) Liu er a/. (1988) Heller er a/. (1984) Lazarenko erol. (1981) Ruocco (1989) Zheng er a/. (1992) Yue e t a / . (1984)

51 38 49 46 49 25 38

50 48 3 2

49 16

~

References

5 6

Rutter er a/. (1990) Forster and Heller (1994)

9.3 E u r a s i a n N e o g e n e

15 1

GPTS LSD El-Dipodomys (A), Smilodon (C), Mammuthus (C)

I Ondatra(A), Synaptornys (A) El Erethizon (A), Neochoerus (A), Glyptotherlum (T), Glossotherlum (T), Stegomastodon (1) Sylvilagus (A) Chasmaporthetes (K), Odocoileus (K), Pratilepus (I) ~ / Equus (I), Trigonictis (I), Castor (I), Ursus (I), Ophiomys (W), Pliopotamys (0, Thomomys (I) m - Pliophenacomys (A,IV0, Geomys (h0 I \ ~ Paenemarmota (M), Prosigmodon (M), Notolagus (M), Taxtdea (M)

HSD El- Stegomastodon (A,C) t Hypolagus (r), Nannippus (T), ~-.iBorophagus (T), Rhynchotherium (T), ( Prodipodomys (A)

II-~ Agriotherium (M), Prosthennops (M), / Dlnohippus (M)

ii/Onohippidium (A)

. . ) Plesiogulo (A), Pliotaxidea (A), II-t Sphenophalus (A) i Calomys (aensonomys) (A) 9 Agriotherium (1"), Rhynchotherium (1), ~-Pediomeryx (1"), Perognathoides (A) Prodipodomys (A), Galushamys (A) 9\~ Plesiogulo (NM), Pliotaxtdea (NM), Dinohippus interpolatus (NM), Astrohippus ansae (NM), Dipoides (NM) ~-Pliometanastas (NM), Epicyon (NM)

Figure 9.2 Loweststratigraphic datum (LSD) and highest stratigraphic datum (HSD) of North American Plio-Pleistocene mammals determined from magnetostratigraphic studies (after L i n d s a y et aL, 1984). L e t t e r s in b r a c k e t s refer to location of fossil finds: A = A r i z o n a , T = Texas, M = Mexico, N M = N e w Mexico, C - California, I = Idaho.

early Hipparion fauna. Sen (1989) has suggested that Hipparion appears in the Mediterranean at approximately 11.5 Ma within Chron C5r. This age, which should correlate to the base of the Vallesian, is the same as the age given by Tedford et al. (1987) for the base of the Clarendonian. Whistler and Burbank (1992) have placed the base of the Clarendonian in C5Ar.3 at 12.8 Ma, however immigrant taxa were not used to define the boundary. If this placement of the boundary is accepted, then the base of the Vallesian is about 1 My younger than the base of the Clarendonian (Fig. 9.3). The base of the Astaracian, the interval preceding Vallesian in the Europe land mammal zonation (Fig. 9.3), is placed by Krijgsman et al. (1994b) in C5ACn at about 13 Ma, much younger than previous estimates, and very different from the base of the Barstovian, which has been correlated to C5Br by MacFadden et al. (1990b). In North America, the boundary between the

| 52

9 Cenozoic Terrestrial Magnetic Stratigraphy

-,i ,IIi

ati]l]m]

i Im ~

ui

i

I

-

mnnmii I

I

<

I

Z a5 Z cr

,~ o.

I I I

-~ll laiR I1 ~

-~11 lllllil

~i i

'

I

II l ~

,

~lillili i ill I

I I

t t

-9

,

~,lilil i /lie

I

I

ue!loJ

n.I.

Je!Selie^ I ue!oeJelsv

ue!ll!qdtueH ,t

ue!uopueJelO

.

,.

I

--

o ii !

:~

~,

,

~ ,

z

-

t

I

ue!^olsJefl '

I

0

~

I I I I I I I I i I I I I I I I I

I

t t

-~l/llgl i i

'

I

~

*

! ic l 14 '

~

"

"

t)

~

ue!ueelJQ

I'pJol6u!tueHI

I

"

"

~ z

mR-

l i i l l c

'

ue!ue6V ueeeJe)l!JV

i

~

1,.

,

!1

I I I

Ill l

I

t

i

i

-~1~--1 !!liB l II !

lllil

I

i

o

I-.

-~1 Ill

I

I

Z

I I I

I I

~ - / ~ / ~ UJ

',

I

~,

I

I I I

II1| ,.-

~

I

,,..

'

I!1 c ,.-

&

'

c

,..

,i,

Figure 9.3 Terrestrial Miocene magnetic stratigraphies. Numbers refer to individual studies

(see Table 9.4).

early and late Barstovian is placed at 14.5 Ma, coincident with severe climatic cooling in the middle Miocene and with the newly determined Orleanian/Astaracian boundary in Spain. Lindsay (1995) has recently placed the base of Barstovian land mammal age to the early part of chron C5Cn.3 (approximately 16.8 Ma) based on the appearance of the immigrant Copemys.

Table 9.4 Miocene Mammals 0

NSE

+35

-1060s -117

IJSA. A n

+34.5

-11358

4

5

C4-CSAB

IJSA. Colo

+3S.3

-118

1

83

3-4

Hepburn's Mesa

CSAD-CSAD

IJSA. Mont

+45.3

-IIl.X

2

62

Verde Fm.

Gauss-C4

IJSA. Ari

+34.S

- I 11.75

9

7 Tesuque Fm 8 Cabriel Basin 9 Kastellios Hill 1 0 Bou Hanifia 11 Aragon 12 Molasse 13 Salt Range

CSA-C5C GilLC4

+36 +39.5 +3S.lX +3h +41.21 +47.57 +32.75

-lM -1.33

7 2

+XI

Mid Miocene SA-SC ChA-C6C Gil-CSA

USA. NM Spain Crete Algeria Spain France Pakistan

+I

1 I

-1.38 +S.S3

2 2

+73

CSB-C6AA

Pakistan

+30.1

CSA-CSN

India

+31.5.5

Rock unit

Lo Age Hi

Chamita Frn Barstow Frn.

GilbertK4A CSAB-CSC

IJSA. NM IJSA. Colo

3 Big Sandy Fm.

GII-C~C

4

Ricardo Gp.

S

6

1

2

Region

A

NSI

NSA

M

D

RM DM

AD

A

1

135

272

3 3

500 965

F A

J

7

B Z-P

F-DI F-V

B

N M Z NCh

%R

RT

22

3 6

36 49

R+ R+

V

6

2

43

R+

F.C.T

0

References

6

IS00

F

-

T

Z-B

V

26

II

39

R+

3

100

F

-

T

Z-V

F-V

12

3

39

R+

MacFadden (1977) MacFadden C I a / (1990h) 6 MacFadden era/. (1979) 7 Loomis and Burbank (1988) 7 Burbank and Barnosky

I64

3

300

K

V

A

B

V

21

5

48

R-

6

3 3

650

I-H

A

6 3

56

R+

S X 49

1

50 Yh 70

-

T

B F-D z-P v z-P v P V Z-P D-I Z-V F-V B F-V

26 10

h

230 42 77 37 357 69 360

+70.75

3

127

3

70(1 - 0

T

z-P

+76.6

1

77

3

950

T

Z-B

4

3

1

A T-A

7 5 F - A

11

10

(lY90)

14

IS

Dera Gazni Khan Haritalyangar

C4A

3-4

210 YO

-

100

2 3

280 27s

3

1700

F -

K -

F

-

T I-X T - T I-J T

-

-

-

r

2 6 4

50 59

-

Bressler and Butler (1978) 7 Barghorn (1981) 6 Opdyke rr a / (1989) 6 Sen er a / . (1986) 5

Sen (1989)

20 127

IS

SO

R+

v

19

7

57

R-

7 Krijgsman e r a / . (1994b) 6 Burbank el a/. (1992a) 8 Opdyke e r a / . (1982) Johnson era/. (1982) 6 Friedman C I a/. (19Y2)

F-V

IS

3

43

R+

5

-

R+

G. J . Johnson

el

a/

(1983)

16 Huch Nala 17 Chinji 18 Khaur

C5-C5AC C4-CSD Ch-CIS

Pakistan Pakistan Pakistan

19 Corral Quemado 20 Ngorora Frn. Ngorora Fm. Honda Gp

GilLC7 CS-CSAA CS-CSAA CSA-CSAA

Argentina Kenya Kenya Bolivia

GilK3A C5 B

Mongolia 'Turkey

Hyargus Nuur Gemerek

+32.65 +32.67 +33.16

+72.2 +72.S +72.52

-27

-67.9

+.6h

+36 135.9

1 2 II

71 IS9 272

3

700

-

3 3

IWW)

2MMl

F F

Y9

3

91 104

2300 800 350 300

K K A K

+.66 -21.95

-65.42

1 2 2 3

106

3 3 3

+4926 +39 13

+93.08 +36.05

2 2 1 X 3 114

3 I

-

J-I -

I

8 0 - - T 170 K -

T T T

B Z-B Z-B

A Z-B T Z-V T Z-B T-A Z-P

T

B P

V

V F~V

I6 33 89

5 II

20

4 4

10

F-V F-V F-V F-V

26

2

V D-l

12 3

3

14 18

4

I

SS

4 44

53 46 44 54

53 59

R+ R+ -

R+ R+ R+ R+ -

5 Johnson el a / (1988) 8 Johnson rr a/. (1985) 7 Tauxe el a/. (1980) Tauxe and Opdyke (1982) 7 Butler el a/. (1984) 6 Tauxe el a/. (1985) 6 Deino er a/. (1990) 9 MacFadden el 01. (19!2oa) 3 Pevzner er a / (19x3) 6 Langereis rr a/. (19x9)

| 54

9 Cenozoic Terrestrial Magnetic Stratigraphy

The magnetic stratigraphy of the Oligocene/Miocene boundary in Europe based on fossil mammals has been reported by Burbank et al. (1992a). The section studied covers 280 m in the lower freshwater molasse in the Haute-Savoie (France) and includes the last two micro-mammal zones of the Paleogene (MP29 and MP30) and the first mammal zone (MN1) of the Neogene in stratigraphic superposition. Unfortunately, a 50-m gap separates probable MP30 faunas from faunas assigned to MN1. The polarity zones are designated R1 to N8 and the gap falls between N4 and N2. The authors, however, choose to place the Oligocene/Miocene boundary at the base of N2, which they correlate to C6Cn.3n. The magnetostratigraphic pattern, however, is ambiguous (Fig. 9.3, column 12). Another possible correlation to the GPTS would be N3 and N4 to C6Bn; and N5, N6, and N7 to C6AA. This would place the MP30/MN1 boundary at the base of C6Cn.ln (Burbank et al., 1992a) and would place the base of the Agenian at this position. Barbera et al. (1994) have placed the Oligocene/Miocene boundary at the base of C6Cn.2n in a section in the Ebro basin of northern Spain. Unquestionably the best studied area of continental Neogene vertebrate-bearing sediments is the Siwaliks of Pakistan and India. These studies were begun in the mid-1970s and have continued to the present. The region is wonderfully suited for magnetic stratigraphy because of well-exposed long sections of very fossiliferous sediments that became magnetized early in their history. In this region it is possible to begin in Brunhes age sediments and systematically work back in time to the onset of sedimentation in the basin (Keller et aL, 1977; Opdyke et al., 1979, 1982; Johnson et al., 1982, 1985; Tauxe and Opdyke, 1982). In many respects the area has served as a laboratory for the application of magnetic stratigraphy to (1) vertebrate evolution and migration (Flynn et aL, 1984), (2) sedimentary processes (Raynolds and Johnson, 1985; Behrensmeyer and Tauxe, 1982; McRae, 1990b), (3) tectonics (Burbank and Raynolds, 1988), and (4) basin development (Cerveny et al., 1988). Almost all of these applications of magnetic stratigraphy to terrestrial sediments were anticipated by N. M. Johnson et al. (1975). In the Siwaliks, the magnetic stratigraphy is now so well known from multiple sections that vertebrate paleontologists usually correlate new fossil finds directly to a nearby magnetic stratigraphy and hence to the GPTS. The correlation in China of the mammal-bearing sequences to the GPTS is known in considerable detail from the Late Miocene to the Pleistocene (Tedford et aL, 1991; Shi, 1994; An et al., 1987, Heller and Liu, 1982, 1984). The importance of the Chinese record since the early Mutuyama is enhanced by the fact that the Chinese loess preserves a climatic record that can be correlated to Plio-Pleistocene climatic cycles (Kukla et al.,

9.4 African and South American Neogene

1988; Kukla and An, have been correlated Other data exist from accessible to Western

155

1989). A mandible and a cranium of H o m o erectus to loess magnetic stratigraphy (An and Ho, 1989). Asiatic Russia but unfortunately they are not easily workers (see Pevzner et al., 1983).

9.4 African and South American Neogene Terrestrial magnetic stratigraphy in Africa stems from the great interest in dating the evolution of man. Long well-exposed fossiliferous sequences are available in central Africa, and magnetostratigraphic data exist for the Gauss to Recent interval in the Omo area (Hillhouse et al., 1986), in the Shungura sequence (Brown et al., 1978), and in Afar (Renn6 et al., 1993). Other studies of importance are by Tauxe et al. (1985) and Deino et al. (1990) in sediments from the Baringo basin correlative to the C5AA to C5N interval. The magnetic stratigraphy of Late Oligocene fossiliferous sediments from the Fayum has been reported by Kappelman et al. (1992). This important study is the first from the terrestrial Paleogene of Africa. Magnetostratigraphic studies on South American mammal-bearing sediments have been concentrated mostly in the Neogene or Upper Paleogene in a series of studies by Marshall et al. (1979, 1982, 1986), Butler et al. (1984), and MacFadden et al. (1983, 1985, 1990a, 1993). South America separated from Africa in the mid-Cretaceous, allowing the South American mammalian faunas to evolve in isolation. During the Neogene, however, North America and South America approached one another and eventually became joined through the isthmus of Panama, allowing the mammals to move from one continent to the other. The timing of the exchanges has been revealed by the magnetostratigraphic dating of sediments that contain the oldest record of the exotic mammal groups. In South America, the oldest of the immigrants from North America, the P r o c y o n i d a e (raccoons), appear at 7 to 7.5 Ma in lower Chron 7 (Butler et al., 1984). In this study, the boundary between Huayquerian and Montehermosan land mammal ages was placed at 6.4 Ma. The second wave of North American immigrants occurs in sediments of the Uquian land mammal age dated by Marshall et al. (1982) from late Gauss to above the Olduvai. In a study of the type Chapadmalal, Orguira (1990) correlated this important migration to the late Gilbert; however, the reverse, normal, reverse (R-N-R) magnetostratigraphic sequence is more likely to correlate to the Kaena and Mammoth subchrons, which would place this migration event in the middle Gauss. In North America, the first appearance of South American forms occurs at the beginning of the Hemphillian; however, the most important migration event took place at the Gauss/Matuyama boundary (Galusha et al., 1984).

| 56

9 CenozoicTerrestrial Magnetic Stratigraphy

It would seem that there is an assymmetry between migrations across the isthmus of Panama, with the most important immigrants such as E q u u s arriving in South America during the middle Gauss, while the South American migrants arrive in North America near the Gauss/Matuyama boundary. These studies illustrate the power of the magnetostratigraphic method in resolving problems of intercontinental faunal exchange. An earlier intrusion of fauna into South America has also been recognized and recorded in sediments of the Deseadan land mammal age. This exotic fauna is of considerable interest since it contains caviomorph rodents and the first South American monkeys. This is a most curious interchange since it is believed that monkeys originated in Africa, and it is not clear how they managed to make the journey to South America. Early Oligocene dates on the Deseadan from South America compounded the problem, because the descendant species in South America appeared to be older than proposed ancestors in Africa. In a study of the Salla beds of Bolivia, MacFadden et al. (1985) were able to redate the Deseadan land mammal age and demonstrate that it was originally dated 10 My too old. They were able to place the first appearance of monkeys at 27 Ma and to demonstrate that rodents appear in South America prior to 28.5 Ma. This study addresses the ancestor problem but does not address the problem of how the animals managed to travel from Africa to South America.

9.5 North American and Eurasian Paleogene Figure 9.4 shows the magnetostratigraphic studies of the Paleogene and Late Cretaceous of North America and Europe ranked in time and identified with author and year of publication (Table 9.5). The boundaries of the land mammal ages in the Paleocene have been updated by studies of Butler and Lindsay (1985) and Butler et al. (1987), and their placement of Paleocene mammalian boundaries is used here. In cases where individual sequences have been restudied and reanalyzed, the most up-to-date reference is given (Table 9.5). Mammalian chronology for the Late Eocene and Oligocene has been reviewed and revised by Swisher and Prothero (1990) and Prothero and Swisher (1992). Our correlation of Duchesnian to Arikareen North American land mammal ages follows Prothero and Swisher (1992) except for some minor revisions. The boundary between the Duchesnian and Chadronian is placed within C17n (Fig. 9.4). This placement satisfies the magnetic stratigraphy and new radiometric results from the Vieja section of west Texas (Testamarte and Gose, 1979). Unfortunately, the magnetic stratigraphy for this section is poorly defined. Prothero and Swisher (1992) placed

| 57

9.5 North American and Eurasian Paleogene

e ~,. 7

2 5 ~.

i~

n

-.-

nr _ _ i . _ .

n__"

_~ 10 r ~ n

-

13

L

A

C

r

A T E E

|

18 n m

D

,

16 [~

i

.

2

"

oo 9 c~

;

17

-m

c

E

-

21 n r

I

-

n 22 r

~

23 n

~

A

n r

~

L

_ -

-

,7,

N

~.

r

24

E "~~ E R

~

14

Y

55-

9

~ n

25 r n 60

27 -

28

n

--

nr~

I

n 29 r ~ 30

P

A L

L

T

E

E ~

0

E

A 9

26

65

;

L 4

50-

2 "",,~

AFRICA

I

n 19 r

45

i~.

II

E -

L ~

" lr

~"

NORTH AMERICA

NEE A RL ~='~

-. 16 rn___...~

-

~-

o__

~

15 n ~ m

40

L

' , =:

3 0 - ~ 12 11 ,nr - - - - .

35-

0

A E R N

'E

I

". 31 n___._~

-'~-

,, m ""

""

_ Jm

5

18

~

L

Y

~. ._~

11

........

Figure 9.4 Terrestrial Paleogene magnetic stratigraphies. Numbers refer to individual studies (see Table 9.5).

Table 9.5 Paleogene Mammals -

Rock unit

Lo Age Hi

Region

A

0

NSE

NSI

NSA

M

83 A 220 A

1 Otay Fm. 2 Flagstaff Rim

Arikreean C12-CI5

USA (CA) USA (WY)

+32.75 -117.25 f42.6 -106.69

2 1

19 114

3 3

3 Washakie Basin 4 Clarks Fork Basin 5 San Juan Basin

C19-C22 C24-C26 C26-C31

USA (WY) USA (WY) USA (NM)

+41.5 +44.8

-109 -109

+36.4

-108

3 3 4

58 273 4

3 3

6 Toadstool Park 7 Vieja Gp.

Cll-C13 C16-CI7

USA (Neb.) USA (Tex.)

-103.63 +30.28 -104.63

1

123

3

1

-

-

-

8 East Fork Basin 9 Black Peak Fm. 10 Dragon Canyon

C20-C22 C24-C26 C27-C29

USA (Wy.) USA (Tex.) USA (Utah)

+43.65 -109.7 +29.40 -103.1 +39.19 -111.3

2 2 2

61 53 35

3 3 3

Hell Creek

C29-C30

USA (Mom.) +47.6

4

39

12 Big Badlands 13 San Diego Area 14 Eureka Sound

Cll-C13

USA (SD) USA (CA) Canada

+43.71 -102.48 +32.75 -117.25 +78.82 -82.45

6 9 3

18

C20-C21 C24-C26

49

C29-C31

Canada

+so.%

- 1 10.72

C12-CI3

USA (Neb.)

+41.9

-103.82

11

15

Red Deer Valley

16 Scott's Bluff

-107

T T

Z-V F-V I

1 8

1 3

T A T

Z-V F-V Z-V V Z-V V

9 7

3 3 3

56 R + 68 47 R +

8 5 7

11 -

4 -

51 -

-

5 1

8 6 6

-

FCT

-

610 K J-X 170 - 110 - 0

T T A

Z-V F-V Z-B F-V Z-B V

7 8 9

2 3 3

74 79 65

R+ R+

3

loo

A

Z-B V

8

2

59

-

101

I 3 3

2000

A-T V Z-V F-V T AT Z-B F-I

27 6 7

3 2 2

66 58 R + 76 R +

4 7 7

1

170

3-4

135 K

A

B

19

3

32

R-

4

1

81

3

135 A

T

-

V

11

3

65

-

4

2-4

152 A I-J

A-T Z-V

I

16

5

49

-

7

3 3 3

150 - J-I 125 - 340 A I

A T T

3 7 18

2 3 4

72 62 59

-

6 5 9

11

4 3 2

43

C18-C22

USA (Tex.)

C26-C27 C12-CI 6 C12-CI6

USA (Mont.) 135.12 -108.95 USA (Wy.) +43.00 -105.5 +29.62 +30.55 Egypt

1

20

3 2

106

Ehro Basin Willwood Fm. Willwood Fm.

C6-C8 C23-C24 C23-C24

Spain USA (Wy.) USA (Wy.)

+41.32 +.21 +44.% -108.42 +44.2 -108.50

1 1 1

108

12

1-1.5 m

53

37 90

3 3 3

-

0

250 K 500 - I-X -

-

-

loo 78

RT

Prothero (1991) Prothero and Swisher (1992) Flynn (1986) Butler era/. (1981a) Butler and Lindsay (1985) Prothero er a/. (1983) Testamarte and case (1979) Flynn (1986) Rapp er a/. (1983) Tomidaand Butler (1980) Archibald era/. (1982) Prothero era/. (1983) Flynn (1986) Tauxe and Clark (1987) Lerhekmo era/ . (1979) Prothero and Swisher (1992) Walton (1992)

-

A-T Z-V V

%R

6 4

AD

_ -

NMZ NCh

References

DM

920 K J-X 2012 - 750 F I - v J-X 225 A K

A

Q

RM

-

17 Devil's Graveyard Fm. 18 Clarks For Basin 19 Dills Ranch 20 Fayum

+29.47 - 103.73

4

D

F-V

Z-V F-V Z-B V Z-P F-V

300 - T Z-P Z-P 1480 A J-Tdl T T-A Z-P 775 A -

F-V F.D-I F-V

8 14

46

-

R+ RR+ R+

6

5 8 7

Butler era/. (1987) Prothero (1985) Kappelman er a[. (1992) Barbera er a/. (1994) Clyde era/. (1994) Tauxe er a/. (1994)

9.5 North American and Eurasian Paleogene

| ~9

the boundary within C16n on the basis of the radiometric dates and an early version of the Cande and Kent (1992a) time scale. The Whitneyan/ Arikareen boundary is placed at the Cll/C10 boundary based on magnetic stratigraphy from the Toadstool Park-Roundtop section in Nebraska (Prothero and Swisher, 1992). The correlation of the magnetic stratigraphies to the GPTS (Fig. 9.4, Table 9.5) is supported in most cases by radiometric ages. An important magnetic stratigraphy from the Devil's Graveyard Fm of west Texas was correlated to C20 (Walton, 1992). A more reasonable correlation, supported by the Bridgerian to Duchesnian fauna, would be to correlate the section to the C21n to C17 interval. One of the problems that can be addressed by magnetic stratigraphy is the heterochroneity (time transgression) of faunal datums. The land mammal ages in North America are based on stage of evolution and appearance of immigrant taxa. Magnetic stratigraphy can test this hypothesis because the evolution and dispersal of mammals and reversals of the geomagnetic field are independent. Two studies in particular have raised the possibility that faunal datums in North America of Paleocene age may be time transgressive. The first of these studies is an extensive investigation of Cretaceous and Paleocene sediments in the San Juan Basin of New Mexico (Butler et al., 1977). This region has a record of the turnover from the dinosaur fauna of the Cretaceous to mammal faunas of Paleogene age. The original interpretation correlated the Torrejonian faunas of the San Juan Basin to C26. The same authors (Butler et al., 1981a) correlated the Tiffanian of the Clarks Fork basin, Wyoming, to C26. This led to the anomalous situation in which two distinctly different mammal faunas, thought to be sequential in time, appeared to be coeval in the western United States. Further study by Butler and Lindsay (1985) resolved the problem by showing that one of the normal polarity zones originally proposed in the San Juan Basin was caused by a viscous magnetic overprint. When this spurious polarity zone was eliminated, agreement was achieved between the two studies. The corrected stratigraphy leads to the placement of the CretaceousTertiary (K/T) boundary in reversely magnetized sediments of C29r. All other K/T boundaries occur in reverse polarity sediments considered to correlate to C29r (Archibald et al., 1982; Lerbekmo et al. (1979), except for a core in New Mexico's Raton Basin which was studied by Payne et al. (1983) and was said to be normally magnetized. Tauxe and Butler (1987) cite a contrary opinion (as a personal communication by E. M. Shoemaker) that the core is entirely reverse polarity. One hopes these data are published so that the results can be properly evaluated. Within

160

9 CenozoicTerrestrial Magnetic Stratigraphy

the resolving power of the paleomagnetic method the K/T boundary appears to be synchronous, between marine and nonmarine sediments. A second study which initially seemed to indicate faunal diachroneity was a study of fauna and flora in Arctic sediments by Hickey et al. (1983). These authors originally claimed that plants and animals appeared in Arctic sediments several million years prior to their appearance in temperate North America. The Eureka Sound Formation, which yielded these results, has been restudied paleomagnetically by Tauxe and Clark (1987), who showed conclusively that the original study was flawed because of the incomplete removal of magnetic overprints and came to the conclusion that no diachroneity is apparent. The results of these studies confirm that land mammal ages are, indeed, useful chronostratigraphic units (Flynn et al., 1984; Clyde et al., 1994; Tauxe et al., 1994).

9.6 Mammal Dispersal in the Northern Hemisphere The fact that mammals disperse from continent to continent is well known, and modern continental faunas are greatly influenced by these processes. The interchange of faunas between the continents has been of great utility in the construction of mammalian biochronologies within the different continents, particularly in North America and Europe. The ability of animals to move from one continent to another is affected by the availability of corridors that allow animals to pass. Plate tectonics has had a first-order effect on the interplay of faunas, since the different plates are in constant motion, changing geographic relationships between continents (McKenna, 1975). The Age of Mammals began at the inception of the Cenozoic, at 65 Ma. Plate tectonic history since that time has been largely one of continental dispersal, which had led to the partial isolation of large continental areas such as Australia and South America. The other continents have a history of intermittent contact and subsequent isolation. Tectonic activity resulting from collision will tend to create links between continental blocks. For example, collision of India with Eurasia during the Cenozoic and the continuing collision of the African plate with Europe provided opportunities for faunal interchange (McKenna, 1975). Faunal dispersal is affected not only by changing paleogeography but also by paleoclimatology, which can erect or destroy ecological boundaries, either expediting or impeding the migration of animals confined to specific habitats. In some cases faunas that appear in midcontinental North America and Europe may have been displaced climatically from the northern part of the continents, as they sometimes have no close relatives in either North America or Eurasia. Paleoclimatology can also create or destroy dispersal

9.6 Mammal Dispersal in the Northern Hemisphere

| ~|

routes by glacioeustatic changes in sea level. These changes can occur rapidly, as has been demonstrated from the marine 6180 record of Pleistocene ice volume change, which implies rapid sea level falls of up to 120 m (Shackleton and Opdyke, 1973). This climatic record can be correlated to the terrestrial faunal record by magnetic stratigraphy. Barry et al. (1985) claimed that evolutionary events and dispersal first appearances in the Siwaliks could be directly related to the 6180 record of the middle Miocene. It is interesting to speculate whether this correlation can be established outside the Indian subcontinent. The change in 6180 for benthic foraminifera from the Atlantic Ocean (Fig. 9.5) (Miller et al., 1991a) shows the overall climatic trend for late Eocene through Miocene time. Mammal migrations have occurred between North America and Eurasia throughout the Cenozoic (Fig. 9.5), and indeed the geochronologic units (land mammal ages) were established using immigrant taxa as part of their definitions (Wood et al., 1941; Woodburne, 1987). It might be expected that the beginning of each new land mammal age would correlate with sea level drops because lowering of sea level would facilitate dispersal. It is unlikely, however, that all dispersals correlate with sea level change, since, in some cases, the corridors might stand so high above sea level as to be unaffected by that change. This is probably the case in the early Eocene when North America and Europe were connected (in the region of the present Norwegian-Greenland Sea), allowing unhampered migration between the two continents. Opdyke (1990) argued that intercontinental mammal migration has resulted from sea level drops following the development of continental ice sheets. Since Neogene land mammal ages are often defined on the basis of immigrant taxa, it was suggested that Neogene land mammal ages in Europe and North America should be closely correlated if both were defined on the basis of immigrant taxa. Miller et al. (1991a) have presented a new and updated 61sO record for the Oligocene and Miocene and argue that continental glaciation began in Antarctica in the latest Eocene (---35 Ma). These authors have formally defined nine oxygen isotope zones in the Oligocene and Miocene which they have designated Oil and Oi2 and Mil through Mi7, ranging in age from the early Oligocene through the early upper Miocene (Fig. 9.5). Miller et al. (1991a) argued that the prePleistocene glacial stages may have been associated with glacioeustatic falls in sea level of 55 to 180 m; however, they argue that the actual maximum drop in sea level in unlikely to have been as large as Pleistocene values of 120m. The fall in sea level postulated above, will, of course, cause the development of unconformities on passive continental margins (Haq et al., 1987; Christie-Blick et al., 1990). Pitman (1978) has argued that the rate of sea level change also controls the position of passive margin unconformities.

] 62

9 Cenozoic Terrestrial Magnetic Stratigraphy

Atlantic Composite

b ~8 0 Ma

...... --

I 9

-

~

3B

I 'e"-- 8

i

4"

--~ - '

10

3A

2.0 I

: 4A 14''- 6

~ M7

15

0

T: '

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I

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~

20 ~)

"-6

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--

1.0 I

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16

9.5 Correlation of mammal migrations to the GPTS (Cande and Kent, 1995) and oxygen isotope events (Miller et al., 1991a). Positions in time of major mammal migrations are indicated by arrows with the number of immigrant taxa.

Figure

Christie-Blick et al. (1990) have suggested that the times of most rapid drop in sea level should coincide with the formation of unconformities. These passive margin unconformities may be detected using seismic reflection data (Vail et aL, 1977; Haq et aL, 1987). Evidence for sea level falls, combined

9.6 Mammal Dispersal in the Northern Hemisphere

| 63

with physical stratigraphy, has been formally organized into sequence stratigraphy. If there is a causal relationship between ice volume changes and sequence stratigraphy, then a correlation should exist between the isotopic record and sequence stratigraphy. Miller et al. (1991a) maintain that such a correlation does, in fact, exist in the Oligocene and Miocene. However, problems in correlating sequence stratigraphy to pelagic isotopic records will remain until the magnetic stratigraphy of continental margin sediments is rigorously documented. Webb and Opdyke (1995) have reexamined the question of migratory events in North America and divide the migratory first appearances into first-order (>8 taxa) and second-order (>5 taxa) migratory events. Cenozoic faunal migration to and from North America and Eurasia is essentially divided into two segments, the first taking place in the Paleocene and Eocene and the second being Miocene and younger. McKenna (1975) and Opdyke (1990) have pointed out that tectonics and paleogeography are the primary controls on mammal migrations whereas climate-driven glacial eustasy is a secondary effect made possible by changes in paleogeography. The large first-order migratory patterns are undoubtedly caused by changes in paleogeography. It is probable that throughout the Paleogene, mammals could pass between North America and Europe over the Thulean land bridge across the Norwegian-Greenland Sea. The route probably remained open until the rifting apart of Svalbard and Greenland, probably at the end of the Eocene (Rowley and Lottes, 1988). The major drop in the temperature of bottom water which occurs in the world ocean at the beginning of the Oligocene (Oil, Fig. 9.5) is not accompanied by any important mammal migration. The reason is undoubtedly that migration across the Thulean land bridge was no longer possible. The other possible route from Eurasia to North America, the Bering land bridge, presented no apparent barrier to mammal migration during the Paleogene. However, at the end of the Mesozoic the Bering region was at paleolatitudes above 80~ almost on the North Pole. The Bering land bridge region has been slowly moving into lower latitudes during the Cenozoic. It seems possible that even though the region was moving toward lower latitudes (70~ the climate was continuing to deteriorate at the end of Eocene and early Oligocene time as indicated by ice-rafted debris around the Antarctic continent. It seems reasonable that the barrier that impeded mammal migration across the Bering land bridge was climatic. By Miocene time, the Bering region probably reached sufficiently low latitudes to allow mammals to cross at times of sea level drop caused by glacial eustacy. The late Eocene "Duchesnian" interchange at about 40 Ma correlates with an unconformity on the continental margin (Miller et al., 1987) and with a major offlap event in the Vail curve (TA 4.1). In addition, this event

| 64

9 CenozoicTerrestrial Magnetic Stratigraphy

has also been correlated with an increase in isotopic values in both planktic and benthic foraminifera (Miller et al., 1987) and with an estimated sea level drop of about 80 m (Prentice and Matthews, 1991). The next major migration event occurs in the middle Oligocene in C9r (R. H. Tedford, personal communication, 1994) where seven new immigrant species first appear coincident with, and following, isotopic event Oi2 (Fig. 9.5). The third major migratory event takes place in the late Arikareen and early Hemingfordian from 21 to 18 Ma. This is the most important period of immigration in the Cenozoic. These migratory events correlate to isotope events Mila and Milb (Miller et al., 1991a). At least four positive oxygen isotope changes in the 18-21 Ma interval suggest that the lows in sea level at this time were comparable to those in the late Pleistocene (Prentice and Matthews, 1991). Two sequence boundaries occur in this interval (Haq et al., 1987), a type 1 sequence boundary at 21 Ma and type 2 sequence boundary at 17.5 Ma, and it is probable that this migratory event is coincident with eustatic sea level fall. The faunal migrations within the late Miocene occur within the Hemphillian LMA. The Hemphillian is characterized by three second-order migratory events at ---9.7 to ---6 Ma, involving 6 or 7 taxa (Tedford et al., 1987). The first of these occurs near the beginning of the Hemphillian in C4Ar.2 at about 9.5 Ma, correlating with 6180 isotope event Mi7 (Wright and Miller, 1992) and with a type 2 sequence boundary in C4A (Haq et al., 1987). The second of the migratory events occurs at 7 Ma, roughly coeval with the Late Miocene carbon shift and a well-defined 6180 event (Hodell et al., 1994). The final migratory event within the Hemphillian is believed to have occurred at about 6 Ma, which is close in time to the position of Gilbert oxygen isotope stages, TG20 and TG22, astronomically dated at 5.7 and 5.8 Ma (Fig. 6.7) (Shackleton et al., 1995b). Lindsay et al. (1984) have placed the Hemphillian/Blancan boundary, which coincides with a first-order migratory event, between the Sidufjall and Nunivak subchrons. Tedford et al. (1987) placed the event at about 4.9 Ma, coincident with isotope stages Si4 and Si6, within the Sidufjall subchron (Fig. 6.7). The final episode of faunal migration is well known to be at, or close to, the Gauss/Matuyama boundary at 2.5 Ma, and is clearly associated with the onset of the northern hemisphere glaciation (Fig. 6.7) (Shackleton et al., 1984). At this time the horse, Equus, migrated to Eurasia and spread rapidly. Lindsay et al. (1980) proposed this correlation from the Siwaliks of Pakistan (Opdyke et al., 1979), and further research in Eurasia has tended to support the appearance of E q u u s at or near the Gauss-Matuyama boundary. Elephas appears in sediments of middle Gauss age in the Siwaliks,

Table 9.6 Terreshial Tectonics Rock unit I 2 3 4 5

Surai Khola Qurhreda del Cura Karrwa Beds Pyrencan foreland Psehauar Bcd\

12 13 14 15

Oliana-Peramola Ricardo Gp. Mecca Hills Morales Fm Rio Azul Siwalik Salla Bed? Manix Basin Contidencc Hills Axhandle Basin

I6

Upper Siwahks

6

7 8 Y 10

II

1.0 Age Hi

Region

A

Nepal Argentina

+2X

CIX-C24

lndla Spain

+34 +42.25

Pleis-PIN

Pakistan

+34

C13-CI8 Miocene Pleis-Plio Plcis-Plio Mio Mio Mio-Oligo Pleis-Plio Pleis-Plio Eo-K C21C33 Pleis-Mio

Spain USA ( C A ) USA ( C A ) lJSA ( C A ) Argentina Pakistan Bolivia USA ( C A ) USA ( C A ) lJSA Utah

+J2.07

+3SW -306') +32.71 -17.17 67.65 -116.5 +35 t3573 - l l 6 . 5 6 +3Y.42 - 1 1 1 67

Pakistan

+33

Mio-Plcis Mio

Pleis-Plio

-30.25

+3S.3 +33.hl

@

NSE

NSI

NSA

M

-8251 hX.5

2

436 112

I 3

5MWI 1150

74.43 - 1.87

Y

21x1

3

XOII

4

300

3

72

7

50

3

-1.32 -1IX -115.9 -119.5 6X.7S 72.46

3 1

IIX)

3 3 I

73.5

1

2 h

I

84 52 49

3

Y

114 216

3

6.5

h

140

I 3

156

275

3 3 3 I 1 3

15

332

3

D

RM

DM

AD

I -

T

-

% - P F-D-I B F-V

XOIM)

F -

-

'r

310

F

-

A - T 7.-B

-

13"J 15lXl F SIX) 2010 2010 K 2lW) F 70 K XO K 1911 K 1050 15lX)

~

F

.I A

K

'1. T I T-A T-A A-T 1 T J - A FA I T-A I F-A -

T

Z-B Z-V

A

N('h

37 26

17

4Y

-

-

h

3

40

Rt

-

5

Appcl e r n / (1491) Beer (1940)

F+ F-

7 h

Burhank andJohnson (IYX3) Burhank er o l (lY92a)

7

V F-V

27 36

3

65

Y

44

R+ Ri-

V

IS

3

hY

R+

-

18

5 h

63 30 XY 63

R?

F+

f'-V V F-D %-P F Z - B F-V B F B F-V 7-P F-V Z-P F-V Z-V V Z-V B Z-P

Z-6

% R R I F<~,'I-. Q

NMZ

V

26

-

X X

F+

7

Raynolds and Johnson (IYX5)

-

7

F+

5 5 24

33 45 76 hh 63

111

5

56

R+

58

-

-

-

R+ R+ R+ R+ R+ R+ R+

13 29 X h

Burhank andTaherkheli (19x5) Burhank er u l (1992~) Burhankand Whistler (1987) Changer a/. (1987) E l l i s r r o l (1993) Jordan e r a / . (1990) McRsr (IYWIa) McRae (1990h) Pluhar er ul. (IYYI) Pluhar Prul. (1992) Talling ef ol. (1994)

7 4 7 7 7 h 6

-

I 2 3 2 2 I I 13

5

Rcfercnccs

-

-

-

| 66

9 Cenozoic Terrestrial Magnetic Stratigraphy

as well as in Eurasia and China (Opdyke et al., 1979; Tedford et aL, 1991). Faunal exchange at this time is very important between South and North America, and the culmination of movement of fauna from South America to North America is correlated to the Gauss-Matuyama boundary in North America (Galusha et al., 1984). Magnetic stratigraphy has proved to be of great utility in dating and correlating the sedimentary record of tectonic events (Table 9.6). One of the best examples is the study of Burbank and Raynolds (1984, 1988) in which they were able to date the tectonic development of the Himalayan

0

1 L \..:.

0

"~

~++++++++++;++++++++++++++m~+++++++++++++++++++++++++++++~i++++++++++++~~++~+ ++++++"+++~+ + ~ .........1 1

2

i

2

m

i

.E_ 4

'

4

i

5

~3

5 ~

6 I,[iii~:,i;~!:,~'~

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System

\

k RR ~.-!

m

3 ~

E9

,A~.'.'..::Grow,ngfolds/ /' /iiiiii~l J 6

Eastem Syntaxis Fluvial System Paleo-Jhelum Fluvial System Paleo-lndus Fluvial System

Figure 9.6 Tectonic and sedimentary history of the Potwar Plateau region (Pakistan) (after Burbank and Raynolds, 1988).

9.6 Mammal Dispersal in the Northern Hemisphere

l ~7

front using the magnetic stratigraphy of the sediments of the Potwar basin (Fig. 9.6). This type of study has expanded to other parts of the world such as the Pyrenees (Burbank et al., 1992b,c) and Argentina (Jordan et al., 1990). Magnetic stratigraphy has also been very important in studies of basin formation and subsidence (Burbank, 1983; Burbank and Johnson, 1983) and for dating of uplift episodes in mountain ranges (Cerveny et al., 1988). Magnetic stratigraphy has also enabled scientists to study the sedimentation process itself and to address problems of stratigraphic completeness (Johnson et al., 1988; McRae, 1990a,b; Badgley and Tauxe, 1990; Friend et al., 1989).

Jurassic-Early Cretaceous GPTS

10.1 OceanicMagnetic Anomaly Record When the Heirtzler et al. (1968) geomagnetic polarity time scale (GPTS) was constructed, it was noted that magnetic anomalies older than anomaly 32 have lower amplitude and are poorly lineated, in both the Atlantic and Pacific Oceans. The lack of lineated magnetic anomalies in this so-called Cretaceous quiet zone implied either that there were no reversals of the Earth's magnetic field for a considerable length of time or that the magnetic properties of the sea floor were such that field reversals were not recorded. The weight of evidence which became available from magnetostratigraphic studies of Cretaceous rocks on land (Ch. 11) demonstrated that the quiet zone in the oceans corresponds to a time of few or no reversals of the geomagnetic field. In the Atlantic, a Mesozoic sequence (M-sequence) of correlatable anomalies prior to this "quiet zone" was identified and named the Keathly sequence by Vogt et al. (1971) (Fig. 10.1). Larson and Chase (1972) succeeded in correlating the Japanese, Hawaiian, and Pheonix Msequence lineations, and Larson and Pitman (1972) were able to correlate these with the Keathley sequence in the Atlantic. The key to this correlation was the recognition of the fact that the Pacific anomalies had been formed in the southern hemisphere and transported into the northern hemisphere by plate motion. These correlations led to the extension of the GPTS into the Middle Jurassic. The anomalies in the mid-Mesozoic M-sequence were initially numbered from M1 to M25. Unlike the Late Cretaceous-Cenozoic part of the GPTS, anomaly numbers were mainly assigned to reverse polarity anomalies (Fig. 10.2); however, the labels M2 and M4 were assigned to normally 168

10.1 Oceanic Magnetic Anomaly Record

]69

KEATHLEY LI NEATIONS Profile latitude ~

(~

.

.

.

lVl - / /

.

.

.

.

.

.,'U l 7 if_

.

.

q" Z ~, t~

Scale at 34~

0 '

200

km

'

27.00

Hawaiian -v ~ ii T -- ''} ] synthetic profile ~ ~ ~ / ~

oScaleat26~ -

'

km

J

M25 M0 Hawaiian block model

Figure 10.1 The Atlantic Keathley lineations correlated to the Pacific(Hawaiian) block model (after Vogt et al., 1971; Larson and Pitman, 1972).

magnetized anomalies, resulting in a rather confusing situation. The magnetic anomaly between M10 and M l l is labeled M10N. This anomaly was originally omitted from M-sequence anomalies (Larson and Pitman, 1972) and was included by Larson and Hilde (1975) (Fig. 10.2) and designated M10N after Fred Naugler, co-chief scientist of the N O A A Western Pacific Geotraverse Project. Jurassic anomalies older than M25 were subsequently identified and the anomaly sequence was extended to M29 (Cande et al., 1978a). Recent accumulation of oceanic magnetic anomaly data in the M0-M29 interval from the Pacific has led to the development of new block models for the Japanese, Hawaiian, and Phoenix lineations (Fig. 10.3) (Nakanishi et al., 1989, 1992; Channell et al., 1995a). Difficulty in correlation and low amplitude of oceanic magnetic anomalies older than M29 led to the concept of a "Jurassic quiet zone"; however, aeromagnetic profiles from the western Pacific have succeeded in identifying lineated anomalies older than anomaly M29 and, as a result, the GPTS has been extended to anomaly M38 (Handschumaker et al., 1988). The magnetic anomaly profiles from the western Pacific imply that the so-called Jurassic quiet zone is an interval of high reversal frequency. These magnetic anomalies are, however, of much lower amplitude than younger magnetic anomalies, and the reason for this remains unclear. The M-sequence marine magnetic anomalies recorded by Larson and Hilde (1975) (Fig. 10.2) in the Hawaiian lineations have provided the means

| 70

10 Jurassic-Early Cretaceous GPTS

HAWAIIAN

LINEATIONS

PROJECTED

0,2

5OO

PROFILES

o GAMMAS i

~

037~

KM

l

300 J

GL32

I I ^ r~

A

NOAA

040 ~

M-25

22

20

17

15 13

11

10 8

4

3 1

0-M

_

SITESITE 166 303 ~,

SITE 105 160 ' s;m-~100 M.Y.

110 . . . .

1410 . . . .

1/0

. . . .

~'1/ SITE 304

.

.

.

.

110 . . . .

1010 MY

PACIFIC M O D E L

Figure 10.2 Magnetic anomaly profiles used to construct the Larson and Hilde (1975) Hawaiian block model (after Larson and Hilde, 1975).

of interpolation between age estimates for most Oxfordian to Aptian time scales. The time scales of Kent and Gradstein (1985) (KG85) and Harland et aL (1990) (GTS89) imply constant spreading rate in the M0-M25 Hawaiian oceanic anomaly block model of Larson and Hilde (1975) (LH75). The KG85 time scale uses the constant spreading rate assumption to interpolate between 119 Ma for the Barremian/Aptian boundary (base CM0) and 156 Ma for the Oxfordian-Kimmeridgian boundary (CM25n). The GTS89 time scale is based on chronogram ages for the Late Jurassic-Early Cretaceous stage boundaries. These ages were used to make a linear recalibration of the KG85 time scale, thereby inheriting the constant spreading rate assumption. The chronograms of GTS89 for this interval are poorly constrained and therefore do not validate the constant spreading rate assumption.

10.2 Numerical Age Control

M2g

'~

M28

17 1

M22,M~ POL7201 M29

POL7201

M2S

M10 M5

Aiiana~ ~/I/l eation /,~set

M17 POL7201 ~ M5 M, ~ 16 / ~ { ~ , M3 Ill

M~9 M !~ /1'~ ^ ~2 A~~ .~ A.~/'~~,1~ M18 ~~'M17~^

nm NNII]BHlll-I ~ll I N N

/~ %1

M16

IN_NNIll I N _ I - J R [

M,11

MO ~ Hawaiian II~jlisneteation

M1 M3,i~ ~~/~l i~n e a t i OIInl MsetSPh~

UI!NNI llll-l[_- ~ i ~

extendedto m29

Figure 10.3 Profiles used to construct the Pacific ( Japanese, Hawaiian, Phoenix) block models in Channell et al. (1995a). The calculated profile is the stimulation from the Larson and Hilde (1975) block model extended to M29.

10.2 Numerical Age Control Since the KG85 and GTS89 time scales were published several numerical estimates for the age of M-sequence chrons have become available. Mahoney et al. (1993) obtained a 4~ age of 122.3 _+ 1 Ma for the basaltic basement of ODP Site 807 on the Ontong Java Plateau. Tarduno et al. (1991) have argued that Ontong Java volcanism is constrained to a short interval in the Early Aptian; therefore this age estimate should be applicable to the Early Aptian. This interpretation was based on the observation that sediments overlying basaltic basement, and volcanoclastic sediments at distal sites, belonged to the Globigerinelloides blowi planktonic foraminiferal zone, which was considered to be confined to the Early Aptian (Sliter, 1992). Lower Cretaceous planktonic foraminiferal biostratigraphy in sections dated with ammonites and magnetic stratigraphy (Coccioni et al., 1992) has, however, indicated that the G. blowi zone extends into the

172

10 Jurassic-Early Cretaceous GPTS

B a r r e m i a n . T h e G. blowi b i o z o n e does not, t h e r e f o r e , restrict O n t o n g J a v a volcanism to the E a r l y A p t i a n . A t Site 878 on M I T G u y o t , the oldest d a t e d s e d i m e n t s (at 25 m m a b o v e b a s e m e n t ) are f r o m the lower part of the C. litterarius nannofossil zone, b a s e d on the first o c c u r r e n c e ( F O ) of R. irregularis and their location b e l o w the " n a n n o c o n i d crisis." Pringle et aL (1993) o b t a i n e d a 4~ age of 123.5 ___ 0.5 M a ( c o n v e r t e d for s t a n d a r d M m h b - 1 at 520.4 M a ) f r o m these basalts. T h e r e m a n e n t m a g n e t i z a t i o n in the basaltic b a s e m e n t indicates a r e v e r s e polarity z o n e overlying a n o r m a l polarity zone. T h e reverse polarity z o n e is p r o b a b l y correlative to CM1.

0-b

a 160~ " 150"

~ " 200 ,

.......

r--

140

Q.} "O

oE 400,

N 130 "1-

.o 600

o~

~" O

._1

t.-

120

"1"

110 100

....

120

I ....

125

I ....

130

I ....

135

I ....

140

I ....

145

I ....

150

800 1000

I

155

....

i .... I ....

i ....

i ....

i ....

~. . . .

i

120 125 130 135 140 145 150 155 Ages (Ma)

Ages (Ma) 0 200 10o

.--. E

-~ -o 200 O E

-~ -o O E

400

o -~ 300

o -~

600

E

.J~

v

E

400

500

o. 800

.... I .... I .... I .... I .... I .... I .... I 120 125 130 135 140 145 150 155 Ages (Ma)

1000 , , , , i , ~ , , i , , , , i , , , , i , , , , i , , , , i , , , , I 120 125 130 135

140

145

150

155

Ages (Ma)

Figure 10.4 The selected radiometric age determinations for M0 (121 Ma), M16-M15

(137 Ma), and M25 (154 Ma) plotted against the (a) Larson and Hilde (1975) block model, (b) Hawaiian block model, (c) Keathley block model, and (d) Japanese blo~zk model. M0 is not present in the Japanese lineations and for (d) the young end of M1 is assigned an age of 123.2 Ma (after Channell et al., 1995a).

10.2 Numerical Age Control

!73

Coleman and Bralower (1993) obtained a U-Pb zircon age of 122 + 0.3 Ma from a bentonite in the Great Valley Group (northern California). This level has been correlated to the C. litterarius zone based on the FO of C. litterarius, however R. irregularis (a more useful marker for the Barremian-Aptian boundary, see Fig. 11.6) has not been found in the section (T. J. Bralower, personal communication, 1993). C. litterarius has been reported from several Barremian sections and the 122 _+ 0.3 Ma bentonite layer from the Great Valley Group may be Barremian. The 4~ age of 124 +_ 1 Ma obtained from reversely magnetized granitic plutons from Quebec (Foland et al., 1986) may be correlative to CM3, or to the shorter duration CM0 or CM1. Ten major igneous complexes with 4~ ages tightly bunched around 124 Ma record only reverse polarity (Foster and Symons, 1979), and it seems likely that the reverse polarity chron is CM3, which has more than three times the duration of

160

M16

150

200 A

E q) t-

t~

r'-

140

4O0 32

:

"13

130

q) : ;

E6oo

i

"4

8

Keathley Hawaiian Phoenix Japanese

m

8O0 x

1000 120

125

130

135

Larson& Hilde (1975)

140

r

145

150

155

120

110

100 160

O B R A D 9 3 (Ma) Figure 1 0 . 5 0 B R A D 9 3

block models.

time scale (Obradovich, 1993)compared to oceanic magneticanomaly

174

10 Jurassic-Early Cretaceous GPTS

any reverse polarity chron in the younger part of the M-sequence. The duration of the interval between CM0 and CM3 is about 3.5 My (Herbert, 1992), leading to an age for the base of CM0 of about 120.5 Ma. A U-Pb zircon age estimate of 137.1 (+ 1.6/-0.6) Ma from the Great Valley Sequence of northern California has been correlated to polarity chron CM16 or CM16n (Bralower et al., 1990). This is an indirect correlation to the polarity time scale using nannofossil stratigraphy. The interval containing the two dated tuff layers was attributed to the Upper Berriasian Assipetra infracretacea subzone based on three nannofossil events: (1) the first occurrence of Cretarhabdus angustiforatus, which was used to define the base of the A. infracretacea subzone, (2) the occurrence of Rhagodiscus nebulosus, and (3) the absence of Percivalia fenestrata. Close inspection of the ranges of the three marker species (Fig. 11.6) suggests that the correlation of the two

-200

160 El

, Kealhley

:

Hawaiian

o,

Phoenix

--

Japanese

150

A

E

v..~

20o

140

G} r

F-,.ql 0

C

"1-

no 40o

130

(i) .,,.&

E

r,D

",4 01 v

_8 60o rn

12o

"~ 3

v

IM17 i r

8OO

1000 120

'

130

.

9

i

i

'

i

......... 1 ~ t 1 2 ...... ! ....

125

:

X

Larson & Hilde (1975)

135

140

GRAD94

145

150

155

110

100 160

(Ma)

Figure 10.6 GRAD94 time scale (Gradstein et al., 1994) compared to oceanic magnetic anomaly block models.

10.30xfordian-Aptian Time Scales

175

volcanic layers to CM16 or CM16n is too restrictive and that the CM16-CM15 interval may be more appropriate (see Channell et al., 1995a). The age of 154 Ma for the young end of CM25 is based on the correlation of this polarity chron boundary to the Kimmeridgian-Oxfordian boundary (Ogg et al., 1984) and to age estimates for this stage boundary of about 154 Ma from California (Schweickert et al., 1984) and Oregon (Pessagno and Blome, 1990). In the California Sierra Nevada, the ammonitebearing Mariposa Formation is a synorogenic flysch supposed to contain the Oxfordian-Kimmeridgian boundary (Imlay, 1961) and is affected by Nevadan orogeny. K-Ar hornblende ages and U-Pb zircon ages on dykes and plutons appear to constrain the age of the Mariposa Formation and the Nevadan orogeny to the 152-159 Ma interval (see Schweickert et al., 1984); however, many of the age determinations are not of high quality by modern standards. More recent 4~ and concordant U/Pb zircon ages from the Klamath Mountains apparently constrain the Oxfordian/Kimmeridgian boundary to the 150-157 Ma interval (see Pessagno and Blome, 1990). The definition of the Oxfordian/Kimmeridgian boundary in the Klamath Mountains is based largely on the overlapping occurrence of the radiolaria Mirifusus and Xiphostylus. This is a highly controversial definition as Mirifusus and Xiphostylus are both present from Aelenian time in Mediterranean sections (see Baumgartner, 1987).

10.3 Oxfordian-Aptian Time Scales The age estimates mentioned above for the base of CM0 (121 Ma), CM16CM15 (137 Ma), and the top CM25 (154-156 Ma) are not consistent with constant spreading rate in LH75 (Fig. 10.4a). Obradovich (1993) used the LH75 block model to interpolate between the same three age estimates: CM0 (121 Ma), CM16/16n (137 Ma), CM25n (156 Ma). The resulting time scale (OBRAD93) implies an abrupt spreading rate change at CM16 in LH75 and other Pacific block models (Fig. 10.5). Gradstein et al. (1994) have presented an integrated time scale (GRAD94) for Triassic to Cretaceous time. For the Late Jurassic-Early Cretaceous interval, constant spreading rate was assumed for most of the LH75 block model with an increase and then decrease in spreading rate in Valanginian and Berriasian time in order to satisfy the three principal age constraints. Synchronous apparent spreading rate changes in different block models may be an artifact of GRAD94 (Fig. 10.6). A new Hawaiian block model (Nakanishi et al., 1989, 1992; Channell et al., 1995a) may represent an improved estimate of a constant spreading rate record for the M0-M29 interval. The three popular ages for CM0,

176

10 Jurassic-Early Cretaceous GPTS

CM16-15, and CM25 are consistent with constant spreading in the new Hawaiian block model (Fig. 10.4b) and the resulting time scale (CENT94) does not imply abrupt or synchronous changes in spreading rate in the Pacific block models (Fig. 10.7). Figure 10.8 illustrates the relationship between CENT94 and other popular time scales. For numerical ages of Late Jurassic/Early Cretaceous polarity chrons and geologic stage boundaries, according to the various time scales, see Tables 10.1 and 10.2.

10.4 Hettangian-Oxfordian Time Scales The Kent and Gradstein (1985) time scale for the Hettangian-Oxfordian intervals was constructed by interpolation between 208 Ma for the base of the Hettangian (Armstrong, 1982) and 156 Ma for the top of the

-200 Keathley

[3 r

Hawaiian

o

Phoenix

&

Japanese

15(1

A

E

140 r"

200

(9 ro

tO

"10 --

(9

M25 400

-r"

13o N o

tt

E

M21

-,4

8

v

120

8OO

i x

1000 '-" 120

125

130

135

:

i

:

:

110

Larson & Hilde (1975)

140

145

150

155

--100 160

CENT94 (Ma) Figure 10.7 CENT94 time scale (Channell et al., 1995a) compared to oceanic magnetic anomaly block models.

10.4 Hettangian-Oxfordian Time Scales

177

160

A

150

v (/)

(1:1 o

oo 140

E L_

c-

0

130

O

KG85

=

GTS89

c,

OBRAD93

&

L~RAD93

120

GRAD94

120

130

140

150

160

CENT94 (Ma) Figure 10.8 Plot of various time scales against CENT94. LH75 (Larson and Hilde, 1975), KG85 (Kent and Gradstein, 1985), GTS89 (Harland et al., 1990), GRAD93 (Gradstein et al., 1995) and GRAD94 (Gradstein et al., 1994), CENT94 (Channell et al., 1995a).

Oxfordian, assuming equal duration of the intervening 50 ammonite zones from Hallam (1975) (Table 10.2). In this interval, the GTS89 time scale (Harland et al., 1990) utilizes the same age tie points (156 Ma for the top of the Oxfordian, and 208 Ma for the base of the Hettangian) and the same strategy of equal duration of ammonite zones; however, the use of a different ammonite zonation (Cope et al., 1980a,b) leads to slightly different stage boundary age estimates. In GTS89, the Bathonian and younger Jurassic stage boundaries were also determined by interpolation from the magnetic anomaly record. The stage boundary age estimates derived assuming equal duration of ammonite zones, from magnetic anomalies, and from chronograms are listed in Table 10.2. For the Hettangian-Oxfordian interval in the time scale of Gradstein et al. (1994), maximum likelihood stage boundary age estimates were derived

178

10 Jurassic-Early Cretaceous GPTS

Table 10.1 Comparison of M Sequence Time Scales Reversed chron

LH75 (Ma)

KG85 (Ma)

GTS89 (Ma)

GRAD93 (Ma)

GRAD94 (Ma)

CENT94 (Ma)

M0 (top) (base) M1

108.19 109.01 112.62 113.14 114.05 116.75 118.03 118.72 118.91 119.06 119.27 119.79 120.21 120.52 120.88 121.49 121.94 122.37 122.82 122.88 123.31 123.34 123.73 124.07 125.10 125.68 125.74 126.22

118.00 118.70 121.81 122.25 123.03 125.36 126.46 127.05 127.21 127.34 127.52 127.97 128.33 128.60 128.91 129.43 129.82 130.19 130.57 130.63 131.00 131.02 131.36 131.65 132.53 133.03 133.08 133.50

124.32 124.88 127.35 127.70 128.32 130.17 131.05 131.51 131.64 131.74 131.89 132.25 132.53 132.75 132.99 133.41 133.72 134.01 134.31 134.36 134.65 134.67 134.94 135.17 135.87 136.27 136.31 136.64

119.15 120.10 122.56 122.88 123.45 125.15 125.96 126.39

120.60 121.00 123.19 123.55 124.05 125.67 126.57 126.91 127.11 127.23 127.49 127.79 128.07 128.34 128.62 128.93 129.25 129.63 129.91 129.95 130.22 130.24 130.49 130.84 131.50 131.71 131.73 131.91

127.16 127.29 127.68 128.62 128.74 128.99 129.41 129.56 129.87 130.41 130.75 131.81 132.63 133.30 135.18 135.94

134.01 134.42 134.75 135.56 135.66 135.88 136.24 136.37 136.64 137.10 137.39 138.30 139.01 139.58 141.20 141.85

137.30 137.37 137.63 138.28 138.36 138.53 138.82 138.92 139.14 139.50 139.73 140.46 141.02 141.47 142.76 143.28

120.38 120.98 123.67 124.04 124.72 126.73 127.68 128.19 128.33 128.44 128.59 128.98 129.29 129.53 129.79 130.24 130.58 130.90 131.23 131.28 131.60 131.62 131.91 132.10 132.70 133.05 133.08 133.37 133.64 133.66 133.93 134.00 134.23 134.79 134.85 135.01 135.25 135.34 135.53 135.84 136.04 136.67 137.21 137.89

M3 M5 M6 M7 M8 M9 M10 M10Nn-1 M10Nn-2 M10N Mll Mll M11An-1 MllA M12.1 M12.2 M12A M13 M14 M15 M16

126.60 126.80 127.25 127.68 127.98 128.32 128.88 129.31 129.71 130.13 130.20 130.60 130.62 130.99 131.26 131.81 132.12 132.15 132.42 132.66 132.68 132.92 132.99 133.19 133.70 133.75 133.90 134.12 134.20 134.37 134.66 134.84 135.41 135.90 136.38 137.72 138.26

140.40

132.40 132.47 132.55 132.76 133.51 133.58 133.73 133.99 134.08 134.27 134.53 134.81 135.57 135.96 136.49 137.85 138.50

10.4 Hettangian-Oxfordian Time Scales

] 79

Table 10.1 continued Reversed chron

LH75 (Ma)

KG85 (Ma)

GTS89 (Ma)

GRAD93 (Ma)

GRAD94 (Ma)

CENT94 (Ma)

M17

136.42 138.16 138.82 139.31 139.46 139.55 140.74 141.28 141.64 141.71 142.47 143.47 144.74 145.29 147.11 147.17 147.23 147.29 147.38 148.36 148.51 148.72 149.15 149.48 149.51 150.24 150.63 151.06 151.09 151.33 151.48 151.79 152.21 152.39 152.73 153.03

142.27 143.76 144.33 144.75 144.88 144.96 145.98 146.44 146.75 146.81 147.47 148.33 149.42 149.89 151.46 151.51 151.56 151.61 151.69 152.53 152.66 152.84 153.21 153.49 153.52 154.15 154.58 154.85 154.88 155.08 155.21 155.48 155.84 156.00 156.29 156.55

143.61 144.80 145.25 145.58 145.69 145.75 146.56 146.93 147.17 147.22 147.75 148.43 149.30 149.67 150.92 150.96 151.00 151.04 151.10 151.77 151.87 152.01 152.31 152.53 152.56 153.06 153.32 153.61 153.64 153.80 153.90 154.11 154.40 154.53 154.76 154.96

138.51 139.86 140.33 140.68 140.80 140.88 141.82 142.25 142.54 142.60 143.21 144.01 145.02 145.46 146.92 146.95 147.01 147.06 147.13 147.91 148.03 148.20 148.80 149.24 149.30 150.31 150.85 151.44 151.49 151.82 152.02 152.46 153.04 153.30 153.52 153.72

140.86 142.51 143.14 143.60 143.72 143.79 144.68 145.08 145.35 145.41 145.99 146.74 147.69 148.11 149.48 149.52 149.57 149.61 149.68 150.42 150.53 150.69 151.09 151.39 151.42 152.10 152.46 152.86 152.89 153.11 153.25 153.54 153.93 154.10 154.31 154.49 155.51 155.69 155.83 156.00 156.14 156.29 156.77 156.85

138.89 140.51 141.22 141.63 141.78 141.88 143.07 143.36 143.77 143.84 144.70 145.52 146.56 147.06 148.57 148.62 148.67 148.72 148.79 149.49 149.72 150.04 150.69 150.91 150.93 151.40 151.72 151.98 152.00 152.15 152.24 152.43 153.13 153.43 154.00 154.31 155.32 155.55 155.80 156.05 156.19 156.51 157.27 157.53

M18 M19n-1 M19 M20n-1 M20 M21 M22n-1 M22n-2 M22 M22A M23 M23 M24 M24 M24A M24B M25 M26 M27 M28 M29

Key: LH75, Larson and Hilde (1975); KG85, Kent and Gradstein (1985); GTS89, Harland et al. (1990); GRAD93, Gradstein et al. (1995); GRAD94, Gradstein et al. (1994); and CENT94, Channell et al. (1995a).

Table 10.2 Jurassic-Early Cretaceous T i e Scales

Stage boundary

Kent and Gradstein (1985) (Ma)

BarremiadAptian HauteriviadBarremian Valanginian/Hauterivian Berriasian/Valanginian TithoniadBerriasian (JIK) Kimmeridgian/Tithonian Oxfordian/Kimmeridgian Callovian/Oxfordian Bathonian/Callovian BajociadBathonian AalenianlBajocian ToarciadAalenian PliensbachianlToarcian Sinemurian/Pliensbachian He ttangian/Sinemurian RhaetiadHettangian (Tr/J)

119 124 131 138 144 152 156 163 169 176 183 187 193 198 204 208

Harland ef al. (1990) Ma Mag. anom. Amm. zones Chronogram

Obradovich (1993) (Ma)

Gradstein ef al. (1994) (Ma)

Channel1 ef al. (1995a) (Ma)

121 127 130 135 142

121.0 127.0 132.0 137.0 144.2 150.7 154.1 159.4 164.4 169.2 176.5 180.1 189.6 195.3 201.9 205.7

121.0 126.0 131.0 135.8 141.6 149.0 154.0

~

124.5 131.8 135.0 140.7 145.6 152.1 154.7 157.1 161.3

166.I

145.O 153.5 156.0 159.8 162.7 166.6 168.2 173.5

178 187 194.5 203.5 208

For Harland ef al. (1990); italics indicate ages adopted in their GTS89 time scale.

125.5 131.0 135.5 137.5 148.0 151.0 156.0 156.3 159.0 159.2 177.0 178.0 182.0 189.2 203.5 210.5

10.4 Hettangian-Oxfordian Time Scales

1131

from selected radiometric ages. The maximum likelihood stage boundary ages were then smoothed by means of a cubic spline fit, with the dependent variable being the unequally spaced maximum likelihood age estimates and the independent variable being the equal duration of ammonite subzones (Table 10.2).

Jurassic and Cretaceous Magnetic Stratigraphy

11.1 CretaceousMagnetic Stratigraphy The majority of Cretaceous magnetostratigraphic studies have been carried out on pelagic limestones in the Mediterranean region, following the pioneering work of Lowrie and Alvarez (1977a,b) on the Gubbio (Bottacione) section in Italy. The principal feature of Cretaceous geomagnetic polarity is the so-called Cretaceous long normal superchron, an interval of prolonged normal polarity lasting about 38 My, beginning just above the Barremian/ Aptian boundary and ending close to the Campanian/Santonian boundary (Fig. 11.1, Table 11.1).

a. K/T Boundary Magnetic Stratigraphy The position of the Cretaceous-Tertiary (K/T) boundary in the GPTS has always been of great interest because of the associated mass extinctions. The position of this boundary in magnetostratigraphic section was first determined at Gubbio (Bottacione) (Alvarez et al., 1977). Since then, the position of this boundary has been accurately determined in more than 14 magnetostratigraphic sections both in outcrop and deep-sea cores. The average position of the K/T boundary in these studies implies a correlation to C29r (0.75) (75% up from base of the reverse chron). However, the location of the boundary within the polarity zone correlative to C29r is not a good estimate of the position of boundary in time due to a substantially reduced sedimentation rate in the Paleocene part of C29r. Cyclostratigraphy 182

11.1 Cretaceous Magnetic Stratigraphy

| I]3

CRETACEOUS

Ma

20 24

22 \

/

'N

Ic3~

5

,/

13

26

74.5

II

84.0 86.3 88.7 93.3

98.5 ~

~

i

~

~

28

Ic3,~

19

112.(

25 ....

i

121

CM1

131.0

-i

135.1

N

12(

CMll

Z

~T]._.~s

m-CMI6

141

. ....

i

_--~.

o~,7

t

N

N

i

Figure 11.1 Summary of some of the more important Cretaceous magnetostratigraphic studies. For key, see Table 11.1.

184 Table 11.1 Cretaceous Rock unit 1

Scaglia

Lo Age Hi

Region

C29N-C34N

Italy. Guhhio Italy. Pogg10 Italy. Gorgo a. Cerhera Italy, Boss0

+43.37 +12.58

1

+14.51 +12.58

1

+43.56 + 12.53

+43.56

Italy, Belluno Italy, Alhian. Contessa Italy Pie’Dosso Italy, Polveno Italy, Fonte Del Giordano Italy Moria

2 Fucoid Mark

34N

3 Maiolica

CMO-CM10

4 Maiolica

CM13-CM19

5

Scaglia

C24-C34

6 Fucoid Mark

C34

7 Maiolica

CMO-CMI

8 Maiolica

CMO-CMI 1

9

Maiolica

CM15-CM19

10 Scaglia Rossa

C29-C34

11 Maiolica

CMO-CMI

12 Maiolica

CMO-CM16

13 Scaglia

C30-C34

14 Fucoid Mark

C34N

Italy Frontale Italy Capriolo Italy Italy Valdorhia

NSI

NSA

M

D

5-1 m

1

250

- J-I

52

3-4

2

.75-1 m

1

+12.53

1

+46.00 +11.75

5

A

Q,

NSE

78

RM

- I

DM

AD

A

A

V

V

NMZ NCh C R RT F.C.T. Q 13

6

22

R+

F+

6

References Lowrie and Alvarez (1977h) Lowrie er a/. (1980h)

A

V

F-V

I

1

-

-

F+

6

180

-

I

T

Z-V F-V

24

II

47

-

F+

7

Lowrie ef a/. (198Oa): Lowrie and Alvarez

.8 m

120

-

-

T

Z-V F-V

17

-

4

Lowrie and Channell

359

127

- -

T

-

5 Channell and Medizza

v

33

7 4 7 13

(1984)

(1984)

28

-

(1981)

30

+43.36 + I 2 5 6

I

72

+45.55 +10.11

3

1.5m

200

+ 10.05

3

1.5m

+43.30 +12.54

1

+43.50 +12.54

Tdl

T-A

Z-P

F-V

15

1

16

R-

-

6

Tarduno era/. (1992)

- I

TA

Z-P

F-V

34

11

41

R+

F+

9

200

- I

TA

Z-P

F-V

34

II

41

R+

F+

9

104

275

-

T

Z-B F-V

21

9

43

-

F+

7

Channell and Erha (1 992) Channell and Erba (1992) Cirilli era/. (1984)

1

250

340

- -

A

V

12

6

18

R+

F+

5

+43.56 + I 2 5 3

1

.75-1 m

1

130

- I

T

Z-V V

15

6 3 3 -

-

6

+45.60 +9.89

I

.3m

I

167

- -

T-A

Z-B V

41

-

5

+43.64 +12.71

1

450

1

300

- I

T-A

Z-V F-V

27

9 2 0 -

-

7

+14.43 112.71

3

24

4

30

- I

A

V

3

1 0 6 -

F+

5

+45.63

-

I

F-V

F-V

15

42

-

Alvarez and Lowrie ( 1978) Lowrie and Alvarez ( 19%) Channell er a/. (1987) Alvarez and Lowrie ( 1984) Lowrie er a/. (198Oh)

+43.56 + 12.53

1

7-1 m 1

+44.38 +4.25 +43.4 - .003.00 +46 +11.76

1 1

163 259

2

+ 46

C28-C32 C34-CM5 C26-C34

Italy Presale France Spain Italy Cismon Italy Cismon S. Atlantic N. Pacific S. Atlantic

Site 463 Site 516 Site 317 Point Loma fm Great Valley Gap Christopher Fm Hole 69OC Cehegin Aix-en-Provence

C34-CMO C29-C34 c34 C32-C33 c33-c34 C34-M 1 C29-C33 CM15-CM17 CM28-CM34

Deccan Traps Sierra Geral

C29-C30 Early Cretaceous Haul- Alb E. EoceneTuronian

15 Maiolica

MGM8

16 Berriasian Lmst. 17 Sopelana 18 Scaglia

CM14-CM19 C29N-C31R C33R-CM8

19 Maiolica

C33R-CM8

20 Hole 525A 21 Site 167 22 Hole 530A 23 24 25 26 27 28

Leg 103 Holes 752-755

94

T

2-V F-V

15

9 4 7 -

-

6

13 X

5 4

Lowrie and Alvarez (1984) Galbrun (1985) Mary er a/. (1991) Channell el a/. (1979) Channell er a/. (1979)

I

1 1 1

394.

T-A Z-V F-V T-A 2-P F-V A B V

2

1

1

394.

A

B

-29.08 +02.99 +07.00 - 177.00 -19.19 -9.39

4

.3-.6 m 88 .35 m

1 1

133 120 630

A AT A

2-B I 2-P F-I 2-B I

Pacific S. Atlantic C. Pacific USA (CA) USA (CA) Canada S. Atlantic Spain France

+21.36 -30.35 -15.2 +32 78 +40.00 + 80 - 65 +38.(!6 +33.5

1 1 1

56

1 I 1 1

IM) 2x9

A A.T. A T A-T T-A A T T

2-P B Z-P Z-V 2-V Z-V 2-V 2-P 2-V

India Brazil

+20

Atlantic Indian Oc.

+11.76

+ I75 -35.18 -1468 - 117.29 -121.5 -92 +2 -1.81 +5.5

+75 -18 - 56 -30 -48 +42.17 -12.17 +93.83 -31

1

I

3 9 6

2 2 33 3 20 6 4

180

74 45 252 76 .25 m 73

I

30 150

1

50 200 3000 IS00 70 12

110

1-4

1166

27 3

4-10 500

loo0

1.5 m 576

383

300

1

310

3 5 I

-

V

F-I I I

R+ R-

-

26

47 50 20

-

-

6 5 6

26

20

-

-

6

-

5 5 4

12 10

II 5 9 3

5 3 8 6 2 4 4 1 5 2 0 1 6 3 5 1 01 2 24 2 4 6 3 05 3 0 7 2 72 6 2 5

V V F I F F

4 13

T.A. 2-P F T-A 2-V F

3 4

2 -

20 26

T-A T-A

2-P I 2-P I

3

3 5 11

-

-

-

-

R+ -

F+

-

F+

R+ -

-

-

-

55 35

R+

-

-

-

11

-

-

19

40

R+

-

Chave (1984) Tarduno er a/. (1989) Keating & HerreroBervera (1984) 5 Tarduno el a/. (1989) 3 Hamilton el a/. (1983) 5 Tarduno (1990) 8 Bannon er a/. (1989) 6 Verosub er a/. (1989) 7 Wynn er a/. (1988) 4 Hamilton (1990) 7 Ogg el a/. (1988) 5 Westphal and Durand (1990) 6 Vandamrne era/. (1991) 7 Ernesto et a/. (1990) 5 7

Ogg (1988) Gee et a/. (1991)

185

| ~6

11 Jurassic and Cretaceous Magnetic Stratigraphy

based on carbonate content, color density, and magnetic susceptibility from South Atlantic cores and Spanish land sections was used to demonstrate that the boundary occurred in time almost exactly in the middle of C29r (Herbert and D'Hondt, 1990; Herbert et al., 1995). Nonmarine terrestrial sediments from the western United States (Butler et al., 1977, 1981b) as well as from Europe (Galbrun et al., 1993) have contributed to the debate concerning the faunal crisis at the end of the Cretaceous by demonstrating that extinctions of marine and nonmarine fauna were essentially synchronous (see Section 9.5).

b. Santonian-Maastrichtian Magnetic Stratigraphy The Late Cretaceous polarity zones at the young end of the Cretaceous long normal interval were recorded and correlated to foraminiferal biostratigraphy in the original Alvarez et al. (1977) study of the Gubbio (Bottacione) section (Fig. 11.2) Subsequent nannofossil biostratigraphy at Gubbio (Monechi and Thierstein, 1985) has provided a correlation of these polarity chrons to nannofossil zones (Fig. 11.3). The correlations of polarity chrons to nannofossil and foraminiferal biozonations at Gubbio have been ratified and refined in the Southern Alps (Channell and Medizza, 1981), Spain (Mary et al., 1991), and in cores from the NW Pacific (Monechi et al., 1985) and South Atlantic (Hamilton et al., 1983; Chave, 1984; Tauxe et al., 1983c; Poore et al., 1984; Keating and Herrero-Bervera, 1984). Due to the scarcity of ammonites in the pelagic limestone sections, polarity chrons are generally correlated to the ammonite-bearing stage stratotype sections through the micropaleontology. Direct correlation of the GPTS to European Upper Cretaceous ammonite zones which define stage boundaries has not been achieved, although the C33r/C33n polarity chron boundary has been correlated to ammonite biostratigraphy in Wyoming (Hicks et aL, 1995).

c. Cretaceous Long Normal From paleomagnetic data available at the time, Helsley and Steiner (1969) postulated the existence of an interval of constant normal polarity from Late Aptian to middle Santonian which could be correlated to long smooth intervals in coeval oceanic magnetic anomaly records (Raft, 1966; Heirtzler et aL, 1968). The presence of this long interval of normal polarity was confirmed by Keating et al. (1975) from deep-sea cores obtained from the Deep Sea Drilling Project (DSDP). Subsequent studies in Italian pelagic limestones resulted in correlation of the top of the quiet zone (base of C34r) to the early Campanian, just above the base of the G l o b o t r u n c a n a elevata foraminiferal zone (Alvarez et al., 1977;

11.1 Cretaceous Magnetic Stratigraphy

] 87

PALEONTOLOGICAL

LITHOLOGY

POLARITY

VGP LATITUDE

-90

0

+90

ZONES

Planktonic ]Calcareous _foraminiferaI nannofossils

Globigerina I = eu_cluDina %,j

STAGE

Pa[eo(ene

Abathompha!us mayaroensls i

MAASTRICHTIAN

GIt. g a n s s e r i

GIt. tricarinata GIt. calcarata

CAMPAN-

GIt. elevata

IAN

r

=-"~

O

,m

GIt. concavata carinata

..{3

SANTONIAN

GIt. concavata concavata

(.9

[ONIACIAN

ql=~

GIt. schneegansi

,,,

TLIRONIAN

GIt. helvetica ~.H. lehmanni rRtl. cushmani

(notzoned)

Planomalina buxtorfi T. breggiensis

o~-,.===

~ "

G. blowi H. similis

Eiffellithus turriseiffeli

_.J _R _cretace~--L _ Parhabdofithus

augustus : Chiastozygus " litterarius

Micranthofithus obtusus

Lithraphidites bollii

limestone

II

normal

marl

~

reverse

~

CENOMANIAN

BIAN APTIAN

BARREM IAN

HAUTERIVIAN

hiatus

Figure 11.2 Summary of Cretaceous magnetostratigraphic studies at Gubbio and Cismon (Italy) (after Lowrie et al., 1980b).

| 88

E

11 Jurassic and Cretaceous Magnetic Stratigraphy

~~ .~ ~

500- ,~ii ~

CALCAREOUS

POLARITY CHRONS

EVENT

44O

top C grandis (484m)

Ii

-

20r

420

•

CPI2

t

P8

lY r base M arag~

--i


(3998m) - - ~

I

~tYmp:ff:Tmi:.ii~top_iRpseudomenardii I-base O. diastypus (386m) "'~pa L _ / top M edgari (382m) "~ ~..base D. multiradiatus (378 5m~,,,j.p r ' ' / t o p lld. velascoensis ( 3 7 6 r n ) \

24r . . . .

I

~a: DI-I~le~nl~ll~3~32?)m) " ' ~ "P, OaseP. pseudomenardii (364~7"/ -~ ~ }

ase F. n is (36 .5 - base E maceltus (354m) --.C~2, -.base C danicus (350.5_m). -'CPI' -top C r e t a c e o u s (347.6m)NC23 - base M. murus (337.9m)

- base L. quadratus ~326m)

NC21 !

- top 77.trifidus (302m)

I 32r

~

'

- base T trifidus (264m)

z z

_ top E. eximius (253m) .- base T gothicus (250m)

33n

I

,

base G. gansserl (302m)

MC$8

Z<X ~__

~

" top G. calarata (2ZO.2m) base G. calarata (264m)

NCl9o

P-"

,r

u~ ,r

' i :~ L

--

~

Z ; 'r

MCs7

" . base B. parco (209m)

NCl8 ~

or} ~

~1

i

~uul

5

~Z

U

- base C. oculeu~ (222m) 33r

-r"

Z

z z

220- .,=1=,=

MCIIO ~

NC19b

--~ ~J ~ ~ I ~ ~. 1 "" ~

base A mayaroensis (330m) ~ base G. contusa (325 5m) --

I~20 MCs9

xz

260- z

I

~,1~=" base M. angulata (359m) .~"Pt--" base M. uncinata (356. 5m) ~ PI -~'base S. trinidaclensis ( 3 5 3 m ) ~ =--"~,base S pseudobuiloicfes (348m~ MCsll base G. eugubina (3476m) /

~lC22 ~ ~

31r

r

z 240. z

base G. palmeroe (418.5m)

CP 9 p 6

31n i

280 9 z r

i

base T. orthostylus (398m)

32n

~

i

CPIO!-- base G. taroubaensis (408 7m) - -.

29r 30n

i

I

base Hantkenina sp. (429m) _ . . _ ~

base D. Iodoensis (402m)

340

~,J

----

iP 9

22r ~ -

360

i

- P1 I

23r ~

25r

300, m =

-

base N fulgens (445m)

i~

~i

base O sublocloensis (417. 5m) base C crassus (414m) ~

26r/ 27r

320- ~ i~

'

,

P1] - - - - - - - P12 o M CP13 ~ t p . aragonensis (453m)

base R. umbilica (463m)

. ~

380

EVENT

CPI~I~ P ~ " base G. semiinvoluta (489m) --.. --

L 21r I

400 "="-=

,

STAGE

CPI4

I

~

~ ZONE, ,

basetrecurvus(50'm) CPl~Pls

'' 16r

9

FORAMINIFERA

ZONE

480 ~i~ 18r/ t z~x~: 19r 460 i

PLANKTONIC

NANNOFOSSILS

! base Giobotruncana elevata (196m);,

Z

NC17 34n .... Figure 1 1 . 3

~ ,

_,1

9

.

.

.

.

.

.

.

.

.

Correlation of Late Cretaceous-Eocene calcareous nannofossil and planktonic foraminiferal events/zones to polarity chrons at Gubbio, Italy (after Monechi and Thierstein, 1985).

11.1 Cretaceous Magnetic Stratigraphy

| ~q

Lowrie and Alvarez, 1977a,b; Channell et al., 1979; Channell and Medizza, 1981) (Fig. 11.2). The entire Cretaceous long normal interval has been recorded in the Cismon section (northern Italy) and here the reverse polarity chron at its base (CM0) is close to the Barremian/Aptian boundary (Channell et al., 1979) (Fig. 11.2). The base of CM0 has since been found to immediately postdate the first occurrence (FO) of nannofossil R. irregularis in two Italian land sections (Capriolo and Pie' del Dosso) (Channell and Erba, 1992) and at ODP Site 641 on the Galicia margin (Ogg, 1988). In the absence of the ammonite Deshayesites, which formally defines the Barremian/Aptian boundary, the FO of R. irregularis is considered the most reliable microfossil marker for this stage boundary. The base of CM0 is usually close to but slightly younger than this nannofossil event. There are a number of magnetostratigraphic records of post-CM0 Aptian-Albian reverse polarity chrons. The first such observation in outcrop was by Lowrie et al. (1980a), who observed a reversely magnetized bed in the Late Aptian Globigerinelloids algerianus zone at Valdorbia (Umbria, central Italy). This so-called ISEA reverse polarity zone (Fig. 11.4), also known as polarity chron C M - 1 (minus 1), was not observed at two other sections in Umbria where the same foraminiferal zone was present in comparable or greater stratigraphic thickness; however, Tarduno et al. (1989) observed two samples with reverse magnetizations just above the FO of G. algerianus at DSDP Site 463. The documentation of this short reverse polarity chron is strengthened by two samples with reverse magnetization spanning 43 cm in a core from ODP Site 765 (Ogg et al., 1991b), where the reverse polarity zone postdates CM0 and is late Aptian in age. Seven reverse polarity zones have been observed by Tarduno et al. (1992) in the middle Albian interval (Fig. 11.4) (A. albianus nannofossil zone and top Ticinella primula to base Biticinella breggiensis foraminiferal zones) at the Contessa section near Gubbio (Umbria, central Italy). In these reverse polarity zones, the magnetization is carried by hematite and the reddening of the sediment may be late diagenetic in origin. The reverse magnetization components in this interval are not antipodal to the normal magnetization components, and the reverse directions are offset toward directions consistent with the Late Cretaceous or Paleogene. The duration of the proposed middle Albian polarity zones is well known from lithologic cyclostratigraphy (see Herbert et al., 1995, and references therein). Of the seven polarity zones recognized by Tarduno et al. (1992), the thickest (3.25 m) represents about 800 ky. This duration is greater than that estimated for reverse chron CM0 and about twice that estimated for CM1 (see Herbert, 1992), both of which are usually recognized by shipboard magnetic anomaly surveys. It is, therefore, doubtful that these middle Albian reverse

| 90

11 Jurassic and Cretaceous Magnetic Stratigraphy

CEN 198"5

PLANKTONIC FORAMINIFERS

CALCAREOUS NANNOFOSSILS

{3)

~=

R. brotzeni

i~

LO R. irregularis L~ F O E . turriseiffefii

m E

OA

EIc

F O P . achlyostaurion FO "small" Eiffellithus

z

!

R. ticinensis ~_ R. subticinensis

~

.o

.J <

FO R. appenninica FO R. ticinensis F O R . subticinensis

T. praeticinensis F O B . breggiensis

T. primula FO A. albianus FO 7. orionatus

H. rischi

E 112.0

F O R . brotzeni

R. appenninica

B

~

.......... OAE1 b ~ H. planispira ..........

FO C. ehrenbergii F O P . columnata

FO T. primula LO T. bejaouaensis

T. bejaouaensis LO M. hoschulzii LO N. steinmannii

Z

N. truittii

H. trocoidea

acme

G. algerianus

<

F O G . algerianus

FO C. achylosum

G. ferreolensis

FO B. africana

L. cabri

9FO E. floralis F O R . angustus FO F. oblongus

121.0

i (N3

rr

FO T. bejaouaensis LO G. algerianus

FO L. cabri

(~Ela

~4-- n

Nannoconid "crisis"

G. blowi ..........

.~:

a~

LO L. cabri

G. duboisi kO C. oblongata

F O G . blowi F O G . duboisi

notzoned

(3~,15

~

~

LO L. bol//i

Figure 11.4 Correlation of mid-Cretaceous calcareous nannofossil and planktonic foraminiferal events to polarity chrons (after Larson et aL, 1993).

polarity zones represent the geomagnetic field at the time of deposition of the sediments. Until the middle Albian reverse polarity chrons are recognized elsewhere, we consider that they should not be incorporated into the geomagnetic polarity time scale (GPTS).

11.1 Cretaceous Magnetic Stratigraphy

|9|

d. Berriasian-Aptian Magnetic Stratigraphy "M-sequence" polarity chrons are numbered according to the correlative oceanic magnetic anomaly. These anomaly numbers, by tradition, correlate to reverse polarity chrons except for M2 and M4, which correlate to a normal polarity chrons. We use a prefix "C" to distinguish polarity chrons from magnetic anomalies and follow Harland et al. (1982) in labeling normal polarity chrons (other than CM2 and CM4) using the number of the next older polarity chron with the appendage "n." Using this nomenclature, CM9 denotes the reverse polarity chron correlative to magnetic anomaly M9 and CM9n denotes the normal polarity chron between CM9 and CMS. It should be noted the magnetic anomaly between M10 and M l l is labeled M10N. This anomaly was originally omitted from M-sequence anomalies (Larson and Pitman, 1972) and was included by Larson and Hilde (1975) and designated M10N after Fred Naugler, co-chief scientist of the NOAA Western Pacific Geotraverse Project. Land section magnetic stratigraphy in the Mediterranean region has been the basis for the correlation of the M-sequence polarity chrons to biozonations and hence to geologic stage boundaries. The oceanic magnetic anomaly record from the Hawaiian lineations (Larson and Hilde, 1975) remains the template for the M-sequence polarity pattern. Some pelagic limestone sections in Italy have recorded substantial portions of the Msequence pattern. Notable among these sections are Gorgo a Cerbara (CM0-CM9) (Lowrie and Alvarez, 1984), Polaveno (CM3-CM16) (Fig. 11.5) (Channell and Erba, 1992; Channell et al., 1995b), Capriolo (CM8CM16) (Channell et al., 1987), and Bosso (CM14-CM19) (Lowrie and Channell, 1984) (see Fig. 11.1, Table 11.1). It is remarkable that the polarity pattern derived by Larson and Hilde (1975) from Hawaiian oceanic magnetic anomalies has been confirmed time and time again by these land section studies, particularly in the CM0-CM19 interval. The only polarity chron observed in land section which is not included in the Larson and Hilde (1975) polarity pattern is a second reverse subchron between C M l l and CM12. These two subchrons were recognized at Capriolo (Italy) (Channell et al., 1987) and have also now been observed in the oceanic magnetic anomaly record (Tamaki and Larson, 1988). The vast majority of the magnetostratigraphic studies in the CM0CM19 interval have been carried out in the Maiolica Limestone Formation of Italy. This is a Tithonian to Aptian white/gray thin bedded magnetitebearing pelagic limestone with fairly constant sedimentation rates (typically --~15 m/My), thereby aiding the recognition of polarity zone patterns. The biostratigraphic control in the Maiolica is mainly from calpionellids for the

Polaveno

Lithology

Nannofossil Eventsand Zones

Polarity Chrons

z ,~ <_

i5o i 'i i--!.ii-'-~.ii 'i.i--.ii 'i..i........ 'i'i~!-.i.'i i .i i .i.il

C M 3 C. oblongata

~

1:

!

en

...... .........

!00

...........................................................

............

L. b o l l i i

CM5

"

CM7 i

.

.

.

.

.

.

i

.

150

R. terebrodentarius

,

9 C. cuvilleri

CM9

t ....

L. bollii

CM'~ON t.

.

N. bucheri

oo ....

T. verenae

~

--

o

8~ t. . . .

CM12 _'1~ - ~ ~

~ . . . . . il ......

. . . .

! .

.

.i. 11... " _ , i - . . .... ~ ~. . . . . . . . . . .. :

i

T. verenae

0

40

<

z

~_

C. oblongata Calpionellites d a r d e d

z


. . . . . . . . .

CM~5

T=



.~

R. fenestratus

t ....... C. angustiforatus

-40

z

~

CM113 I

t

-80

o

~ .

80

VGP Latitude (o) Figure 11.5 Virtual geomagnetic polar (VGP) latitudes from the Polaveno section (Italy), correlation to the GPTS, and correlation to nannofossil and calpionellid events (after Channell and Erba, 1992; Channell e t al., 1995b). In the lithologic log, F indicates small faults (with minimal offset), speckles indicate cherty intervals, diagonals indicate marly interval, and black bars indicate black shales. Horizontal thin bars indicate pelagic limestone, the dominant lithology.

11.1 Cretaceous Magnetic Stratigraphy

| 9~

Tithonian and Berriasian and from calcareous nannofossils for Tithonian tq Aptian. It has been demonstrated in numerous Maiolica sections that polarity chrons correlate consistently to nannofossil events (Fig. 11.6) and calpionellid events (Fig. 11.7). For correlations to calpionellids, see Channell and Grandesso (1987) and Ogg et al. (1991c); for correlations to nannofossils see Bralower (1987), Channell et al. (1987, 1993), Bralower et al. (1989), Ogg et al. (1991b), and Channell and Erba (1992). The result of these studies is that the combination of nannofossil or calpionellid biostratigraphy and magnetic stratigraphy gives a very precise and useful integrated stratigraphic tool for pelagic limestones of this age. Even for short sections where polarity zone patterns are not distinctive, the nannofossil/calpionellid events indicate the approximate stratigraphic position relative to the GPTS, and the magnetic stratigraphy then gives precise stratigraphic control. Although the correlation of polarity chrons to nannofossils and calpionellids is firmly established, the correlation to stage boundaries requires correlation to ammonite zones. Although ammonites are very rare in the Maiolica limestones, recent finds place the Hauterivian/Barremian boundary in the upper (younger) part of CM4 (Cecca et al., 1994) and the Valanginian/ Hauterivian boundary at the young end of C M l l (Channell et al., 1995b) (Fig. 11.6). Previous estimates of the correlation of these stage boundaries to polarity chrons relied on the correlation of nannofossil events to ammonite zones from Thierstein (1973, 1976). For the Berriasian-Valanginian boundary, the relevant ammonite zonal boundary has been correlated directly to CM15n at Cehegin (Spain) (Ogg et al., 1988). For the Jurassic/ Cretaceous (Tithonian/Berriasian) boundary, there are two alternative definitions with respect to ammonite zones, with no clear consensus. In the Tethyan realm, the boundary lies either at the B. grandis/B. ]acobi zonal boundary or at the underlying B. jacobi/Durangites zonal boundary. At Carcabuey (Spain), the B. jacobi/Durangites zonal boundary lies in a normal polarity chron at the top of the section, which is interpreted as CM19n (Ogg et al., 1984). Elsewhere, nannofossil and calpionellid events in the vicinity of the Jurassic/Cretaceous boundary have been consistently correlated to polarity chrons in land sections and deep-sea cores (Lowrie and Channell, 1984; Cirilli et al., 1984; Channell and Grandesso, 1987; Bralower et al., 1989; Ogg et al., 1991c). These studies imply that the Jurassic/Cretaceous boundary lies in the upper part of CM19n or at the CM19n/CM18 polarity chron boundary, and it has been advocated that this polarity chron boundary be used to define the stage boundary (Ogg and Lowrie, 1986). At Berrias (France), the Berriasian stratotype section yielded a magnetic stratigraphy that is correlated to calpionellids, nannofossils, and ammonite zones (Galbrun, 1985; Bralower et al., 1989); however, the base of the section is within the B. grandis zone and

194

11 Jurassic and Cretaceous Magnetic Stratigraphy

POLARITY CHRON

Ma

. .1 2.o .

CALCAREOUS NANNOFOSSILS

.,

"--" -

CM~

CM1

-

--

Nannoconid crisis ~

---r

[]~

~ F. oblongus R. irregularis

~

125 CM3

-- CM4 "-'~r--

AMMONITES POLARITY ZONES SUBZONES STAGES CHRON

m

]-"1C. oblongata ]--1L. be/Ill

~Rterebrodentarlus 7C." cuvillieri -1 CM10 i "j]-J L. boIlii

D. weissi ~ D. tuakyricus _ ~ - C.'s,~ a'ra"si_ ~_ il. airaudi ,(i) H. feraudi ca. H. sartusi = , A.vandenheckei H. caillaudi ._ "S, nicklesi , ~ S. hu(:Jii ' .9o P.ancjulicostata P. catulloi P.angulicoit. B. balearis- P.'li,qatus" e"

"-t--c I-- 8

--

CM1 -CM11A - ~ - CM12

,

'j_.j T. verenae

~

--

--

CM1

51,.~

C. Ioryi A. radiates

N. pachvdicranus

'

~_j p. fenestrata

T. pertransiens otopeta

m7uR. nebulosusL..F, boissieri _ CM16 j_Jc. angustiforatus 0_.t = - ' ' - ' - ' -

1

J

CM17

~

~

CM18 ~

N. St. steinmannii i " C. cuvillieri ~ ~' "I-JN._ stoinm, minor

Ii

mmll

C. mex. minor

t1" ,

<E rn

m z .

I

or ~

M~

CM1

CM4 C~M7

CM10

I

z~: i ~ 9 ~

UM11A 9 , CM 12

X .r.

i

alpillensi$

B. picteti ~ m - ~ - i.paramimouna ~" (/) ~-- , CM16

T. occitanica

D. dalmasi B. privasensis

<{: ~

T.subalpina i ~

i

io m, ~

B~aco~l

Durangites

~

~=

P. transitoris

Simplisphinctes

B. peroni

g. aa'mTr~n d-um /

R. richteri

,

~--CM10N ~

_ B. grandis. . . . . . .

--' ' ~ - ' ~ l N. colomiin i-1 i, CM19 ~ "J R./aline,~ L_LjI -U. gr. granulosa ' 7J. ...J carniolensis i JM.L.chiastius -~J C. mex. mexicana 17 145 c a 2 0 ~ ----'-I _ P. beckmannfl L...u CM21 ~ j ~_j P. embergeri CM22

H. trinoclosurn

S. verrucosum B. campylotoxus

3S~TM~4"mt'j-JC. oblongata

l

z ~ ,,,

o. =

L. nodosoplicatum

_

<

o.~~~_. ~. ~ ~ T< 'r

S. sayni

.CMlON=~J---iN'bucheril T. verenae

z.

"~ ~ ~)

/ S . biruncinatum

:~

~ E

CM17

CM18

i

CM19 CM20 1 CM21

V~all~ed[n~m:

H. hybonotum

CM22

kl "

"

x

Figure 11.6 Summary of correlation of Oxfordian to Aptian polarity chrons to nannofossil events/zones and ammonite zones. Open bars indicate the range of the nannofossil events with respect to polarity chrons (from Channell et al., 1995a).

11.1 Cretaceous Magnetic Stratigraphy

] 95

~mmonite Calpionellid Zone zone i event ~ P ~

Stage

Busnardoiles

>

T" P ~ r t r a n s i e n s

]

I7ili!ii7i7777!i,~~C.

T. otopeta -_......................

9

B. callisto

da~ri

CM14

=

~ h

D

P. pecteti

" ung;'rica

CM15

:i:!:i:iiii

Z<~ .paramimounum ........B................. ~:~....:~:r~C ~ oblon~ CM 16 I--'C.

h'q <

O. d a l m u i

::::::::::::::::::::::::::::::::J .

::::::::::::::::::::::::::::::::::::::::::::::::: B. privasensis ................................. :!:i:i:7:7:7:i:~:7:7:7:7:7:7:7:7:7:7:7:7:7:7:7:7:! :i:!:i:i:i:i:i:i:i:i:!:!:!:{:i:i

.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:

~

,~'IC. elliplica

B. grandis .......

simplex

C

. ferasini

:ii!i!i!i!iiiiiii!i!i!i!i!iiiiiiiiiiiii!iiiiiiiiii::::::::::::::::::::: i:i:i:!:i:!:!:!:i:::::::::::::::::::::::::::::::: :::::::::::::::::::::::::::::::: ....._ ....... ._.~:::::::::::::::::::::::::::: :, .

.

.

.

.

.

.

i

.

CM18

-

!iii!iiiiiii!i!i!i!!i!!!i!iiiiii!

Z

........o:.,._..~,,,:.......

~)

Sid~l,s;Uincteg::i!::::::::::i:i i i :2i!i;~l.

I-

S ~.;~o;'.:;umChit" I ~,

A

I c-='"n~ ~

,,

:::::::::::::::~.:E:::::-::::::::::::............... T. L'arpathicaT'T B. peroni

s. biruncinatum R richteri

( ~ 20

C. bone. t~,.

~

~I~Chitinoidell,

(3vi21

BB

H. verruciferum

Figure11.7 Correlation of calpionellid events to polarity chrons (after Channell and Grandesso, 1987; Ogg et aL, 1991c).

therefore the section does not include either of the two definitions of the Jurassic/Cretaceous boundary. Ogg et al. (1991c) have identified CM17 in the Purbeck Limestone of southern Britain, providing a first M-sequence magnetostratigraphic correlation out of the Tethyan into the Boreal Realm.

196

11 Jurassic and Cretaceous Magnetic Stratigraphy

11.2 JurassicMagnetic Stratigraphy a. Kimmeridgian-Tithonian Magnetic Stratigraphy As for the Early Cretaceous (CM0-CM18), the Larson and Hilde (1975) oceanic magnetic anomaly block model is the template for geomagnetic polarity in the Tithonian and Kimmeridgian stages of the Jurassic (CM18CM25). Subsequent extension of the oceanic anomaly block model to M29 (Cande et al., 1978), and to M38 (Handschumacher et al., 1988) has extended the record into the middle Jurassic interval. Land section magnetic stratigraphies have been correlated to CM18-CM25; however, correlations to older M-sequence chrons and between-section pre-Kimmeridgian correlations have not been adequately achieved (Fig. 11.8, Table 11.2). The problem for pre-Kimmeridgian magnetic stratigraphy is twofold. First, preKimmeridgian calcareous nannofossil biostratigraphy does not yet allow precise correlation among Tethyan sections; and second, pre-Kimmeridgian Jurassic facies in the Tethyan realm are often highly condensed and/or highly siliceous. The oldest polarity chron recorded by the Maiolica Limestones of Italy is CM19. The Maiolica limestones are usually underlain either by siliceous limestones (Calcari Diasprigni) or by condensed nodular limestones (Ammonitico Rosso). The Sierra Gorda and Carcabuey sections of southern Spain are the key sections for the correlation of CM19-CM25 to ammonite zones (Ogg et al., 1984). In these sections, the "Ammonitico Rosso"-type sediments have mean sedimentation rates of 2-3 m/My. The polarity patterns are distorted by variable sedimentation rates and it is not easy to correlate polarity zones to polarity chrons. Nonetheless, the correlations of Ogg et al. (1984) place the H. beckeri/H, hybonatum ammonite zonal boundary which defines the Kimmeridgian/Tithonian boundary in CM23n, and the I. planula/S, platynota zonal boundary which defines the Oxfordian/ Kimmeridgian boundary in CM25 (Fig. 11.6). Ammonites are lacking in this interval in the Belluno Basin and Trento Plateau (Southern Alps, Italy), where the Oxfordian/Kimmeridgian and Kimmeridgian/Tithonian boundaries lie in the unsubdivided "Saccocoma Zone." The lack of calpionellid or other microfossil events in the Italian sections precludes the accurate definition of these boundaries; however, the polarity zone pattern in these sections is more readable than in the Spanish sections. Estimates of

Figure 11.8 Summaryof some of the more important Jurassic magnetostratigraphic studies. For key, see Table 11.2.

JURASSIC 14

Composite (Ma) ~---144.2 r cO c-

10

54.1

mm

C

EE m

,.,..

0 X

,-m

0 --159.4 c"

m mm

> 0

64.4 r-

6

11

tor

--169.2 C o 0

n~

-s ---<180.1 c o~

12

15

17

.0

.d

0O r .m

m

m m

II

13"195.3 E C

II

03 --.201.9

7"

198 Table 11.2 Jurassic Rock unit

Lo Age Hi

Region

A

@

NSE

NSI

NSA

M

D

RM

DM

AD

A

NMZ NCh %R RT

F.C.T

1

13

-

-

T

Z-P

F-D-I

21

-

60

R-

- I

T-A Z-V F-D-I

15

-

SO

R+ -

6

Q

References

1 Marine Magnetic

Anomalies 2 Sierra Harana

Spain

+37.2

-3.7

1

3 Toarcian

Bathoniard Bajocian Toarcian

France

+47.W

-.02

2

.I m

1

5

4 Carcabuey 5 Ammonitico-Rosso

CM18-CM25 Ox-Callo.

Spain Italy

+37.5

-3.3 +I1

2 3

.I m .05 m

1

+46

1

10 10

- - -

T T

Z-V FV Z-P V

28 55

-

4 0 R + 32 - -

6 La Fuente

Bathe.-Bajo.

Spain

+37.5

-3.3

2

1

10

- -

T

Z-P

19

-

42

R + F+

7 Breggia

Baj.-Car.

Switzerland

+45.87 + 9

2

I

120

- I

T-A Z-V F-V

82

-

47

-

USA Italy

+40.25 -75.25 +45.7 f11.46

3 1

4 139

5 1

4180 60

T Z-P FV T-A Z-B V

1

- -

25

1 9

0 52

-

105

1

13

- -

T

Z-P

F-D-I

22

4

54

R-

F+

-

-

T

Z-P

F-D-I

51

-

41

-

F+

T-A Z-V F-V

82

-

47

-

-

- -

T

D-l

14

-

62

-

C+

- I

T-A -

-

17

6

4 9 - -

11

-

55

8 Newark Supergroup Het-Carnian 9 Xausa CM14-CM23

106

9

9

.2 m

10 Aquilon

L,Oxfordian M Spain

+41.3

-l.W

5

11 Carcabuey

Bath-Aal

Spain

+37.5

-3.3

2

.I5 m

1

33

12 A l p Turati

Ba-Car.

Switz

+45.87

+9

2

.2 m

1

120

1-13.5

3

169

I

24

3 9 0

1

13 Kandelbach Grab.

Sin:Het.

Austria

+4.7

14 Foza

Ber.-Kim

Italy

+45.54 +13.5

Hungary

+47.1

15 Bakonyncoke

Pliens

+I8

1

.I m

1

-

K

I

- I

8.9 - I

T

B

F-D-I

Z-V V

6

-

R + F+ F+

R-

-

Steiner era/. (1989) 7 Galbrun ef a/. (1988a) 6 Ogg er a/. (1984) 6 Channel1 er a/. (1990a) 8 Steiner ef a/. (1987) 7 Homer and Heller (1983) 9 Witte era/. (1991) 5 Channell era/. (1987) 7 Steiner er a/. (1985) 8 Steiner er a/. (1987) 7 Homer and Heller (1983) 6 Steiner and Ogg (1988) 2 Channell era/. (1982b) 5 Marton er a/. (1980)

16 Quero

Titho

Italy

17 Cingoli

Pleins-Sinme

Italy

+43.33

18 Frisoni

ritho-Kimm

Italy

+4.5.53

Ammonitico-Rosso

L Toarcian E

Spain

+ 37.39 -3.49

Kayenta Fm.

Pliensbachian

USA

+38.6

-109.6

101

Krakow Uplands La Luna

M.Ox. E.. Call. Sant. Sen

Poland Veneiuela

+SO.I +Y.Y

+I9 6 -61.0

204 48

1

Lebombo Grp. Morrison Fm.

E. Jurassic Kim-Ox

S. Africa USA

-24 +3X.13

+31.75 -108.21

25 215

4s 1

6000

Morrison Fm

Kim-Tith

IJSA

+MI

-10x.z

I

1

Purbeck Ls. Sierra Gerral

Ber.-Tith

England

+SIM

-2.24

I

I

Brazil

-

-

Sierra Paloma

Toarcian

Spain

+40.65

- 1.1

Summerville & Curtis Fm. Umbria

Callovian

USA

+3x.x

-111.1

TithonianToar.

Italy

+43.33

+I3

I

74

+I321

1

64

20

+ 11.33

I

132

xn

1

.I3

I

1

50

1

20

1

I00

I

.3 7x ~

~

T

Z-B

T

V

6

4

39

6

Z-V F-V

16

-

5x

6

T

Z-B V

IX

7

47

6

T

%-P F ~ D - I 21

53

R

6

T

-V

IY

R+ -

5 7 7

D-l

8 ~

R+ -

A

63

~~

4

30 7x

T Z-P T-A Z-P

F F-D-I

A-T B T V

F D-I

165

T

Z-V

F-D-I

XI1

1I15 -

T

V

F-D-I

33

I07 50

8.5

I

K

A

~

~

19

5

-~~

6 4

R+ -

-

4

5

~ ~

Channell and Grandcsso ( 1987) Channell el a / . (19@) Channell and Grandcsso (1987) Galhrun er a1 (1990) Steiner and Helslcy (1974)

Ogg e r n / . (19Yla) Ca.itillo er a / . (1991) Henthorn (1981) Steincr and Helsley ( 1 9 7 5 ) Strmer and Helslcy (197Sb) O g g r r a l . (1991~) Valencio er a / . (1983)

47

1

75

T

Z-V V

4s7

I

120

T

V

1

2IN

T

Z-B F-V

5

I

F

14

3h

4

Y

92

5

Galhrun er a1 (IYXXh) Steiner (197Xi

14

14

h

Channell (19x4)

YI

ul.

199

~00

11 Jurassic and Cretaceous Magnetic Stratigraphy

the location of the Kimmeridgian/Tithonian boundary in the Italian sections place it in CM22 (Ogg et al., 1984; Channell and Grandesso, 1987). Early work on Jurassic magnetic stratigraphy was carried out on the Kimmeridgian-Oxfordian Morrison Formation in Colorado (Steiner and Helsley, 1975a) and the Callovian Summerville and Curtis formations in Utah (Steiner, 1978). Correlation to European sections is hindered by poor biostratigraphy in these terrestrial and near-shore sediments, complex magnetic behavior which compromises the fidelity of the magnetostratigraphic records, and variable sedimentation rates which distort the polarity zone pattern. Steiner et al. (1994) have made the case that magnetic stratigraphy can be used as a means of correlation in the Morrison Formation in Colorado and New Mexico.

b. Oxfordian-Callovian Magnetic Stratigraphy The correlations of European land section magnetic stratigraphies to M0M25 oceanic magnetic anomalies are fairly robust; however, magnetostratigraphic correlation to M26-M38 has not been well established. CM26 to CM30 have been correlated to Oxfordian ammonite zones in northern Spain (Steiner et al., 1985/1986; Ju~irez et al., 1994, 1995); however, the correlation between land sections and oceanic anomaly records remains somewhat tenuous. The difficulty in correlation is partly a result of discontinuous and low mean sedimentation rates (---1-3 m/My) in the condensed limestone facies. Similarly condensed Callovian-Oxfordian sediments in Monti Lessini (northern Italy) yield polarity reversals and sporadic ammonite control (Channell et al., 1990a) but the polarity pattern cannot be correlated to the oceanic magnetic anomaly record. The existence of polarity reversals in the Callovian-Oxfordian is confirmed by studies of short sections from the Krakow Uplands (Poland) (Ogg et aL, 1991a); however, here again the lack of long, continuously deposited sedimentary sections does not allow a convincing correlation to the oceanic anomaly record. Although the magnetostratigraphic correlation to the M26-M38 oceanic magnetic anomaly record has not been achieved, it is clear that the Jurassic "quiet zone" in the central Atlantic magnetic anomaly record is not due to a prolonged interval of normal polarity, as had been implied by magnetostratigraphic study of the Valdorbia section (central Italy) (Channell et al., 1984). It is now clear that the Callovian-Oxfordian interval, which correlates to the Jurassic quiet zone, was an interval of frequent polarity reversal (Fig. 11.8, Table 11.2).

c. Pre-Caliovian Jurassic Magnetic Stratigraphy For the Bajocian and Bathonian, several sections from southern Spain indicate a very high frequency of polarity reversal (Steiner et al., 1987), an

11.2

Jurassic Magnetic Stratigraphy

20|

average reversal rate of at least 5.5 reversals/My for the Bajocian. The mean sedimentation rate in these sections is in the 1-4 m/My range. The high frequency of reversal and the lack of any discernable "fingerprint" in the polarity zone patterns have precluded clear correlation among land sections. The elucidation of the magnetic stratigraphy in this interval will be aided by studies of more expanded sections. Such sections occur in central Italy but they lack ammonites and therefore await refinement of nannofossil and radiolarian biostratigraphies in this interval. One of the most important magnetostratigraphic studies in the Jurassic is that of Horner and Heller (1983) from the Breggia section (southern Switzerland) (Fig. 11.9). The sedimentation rates in the nodular limestones at Breggia are about 14 m/My for the Pliensbachian and 4-7 m/My of the Toarcian and Aalenian. These sedimentation rates are several times greater than for more typical Ammonitico Rosso-type limestones, and this results in a considerable improvement in the clarity of the magnetostratigraphic record. In addition, the ammonite biostratigraphy for the Pliensbachian-early Bajocian interval at Breggia has been determined in detail by Wiedenmayer (1980). Mfirton et al. (1980) acquired a magnetic stratigraphy from the VGP LATITUDE .90 ~ 0o +90oP~

AMMONITE ZONES

A M MONI TE ZONES

VGP LATITUDE Polarity

Stage

-90

(Ma)

~

0 ~

Stage

+90 ~

,~ (Ma)

sowerbyi oo m

margari-z__.

220

176.5

tatus

~

192

murchi- ~"' ..J

sonae

_

davoei

,~

opalinum

E2

~ 2oo g

180.1

~5

meneghinii

ibex

~o ~8o

i

8

erbaense

bifrons

falciferum z

%1

,.me.on,

189.6

spinatum ~ 0 a

Figure 11.9 Carixian (Pliensbachian) to Bajocian magnetic stratigraphy and ammonite zones at Breggia (Switzerland) (after Horner and Heller, 1983).

202

11 Jurassic and Cretaceous Magnetic Stratigraphy

9-m-thick Bakonycsernye section (Hungary). Here sedimentation rates in the Pliensbachian are about 1.5 m/My and 10 reversals are recognized in this stage. Pliensbachian sections in Italy with similar sedimentation rates yielded comparable numbers of reversals (Channell et aL, 1984). However, at Breggia, the sedimentation rate for the Pliensbachian is an order of magnitude greater (---14 m/My) and 39 reversals are recognized in the stage. Correlation between Breggia and Bakonycsernye shows that seven normal polarity zones at Breggia have been concatenated into one at Bakonycsernye. There is clearly a problem in resolving complete magnetic stratigraphies in the Ammonitico Rosso-type facies when sedimentation rates are a few meters/My, probably because of frequent lacunae even in sections where ammonite finds imply continuity of sedimentation. The Toarcian stratotype sections at Thouars and Airvault (France) are gray marls and limestones, and their magnetic stratigraphies have been resolved by Galbrun et al. (1988a). As the sections are about 5 m thick and the entire stage is represented, the mean sedimentation rates are less than 1 m/My. Nonetheless, a reasonable magnetostratigraphic correlation can be made from the stratotype sections to the Breggia section. The magnetic stratigraphy provides a means of correlation of the West European ammonite zonation in the stratotype sections to the Tethyan ammonite zonation at Breggia. For the Hettangian and Sinemurian, three red nodular limestone sections in Austria (at Kendelbach Graben and Adnet) have been studied by Steiner and Ogg (1988). Correlations among sections are hampered by the condensed nature of the sedimentation and the very poor biostratigraphic control in these sediments. The same problems affect gray/white pelagic limestones of this age in central Italy (Channell et aL, 1984). The best quality Hettangian/Sinemurian magnetic stratigraphy is from a core drilled in the Paris Basin (Yang et aL, 1996) indicating high reversal frequency in the vicinity of the Hettangian-Sinemurian stage boundary.

11.3 Correlation of Late Jurassic-Cretaceous Stage Boundaries to the GPTS Late Jurassic and Cretaceous chronostratigraphy is based on stage stratotypes defined by ammonite zones. However, due to absence or uneven distribution of ammonite faunas in land sections and oceanic cores, Cretaceous chronostratigraphy is often based on calcareous microplankton biostratigraphy and the supposed correlation of these events to ammonites. A review of calcareous nannofossil events in sections with ammonites, microplankton, or magnetic stratigraphy has revealed the uncertainties in correlation of nannofossil and calpionellid datum planes to stage boundaries (Figs.

11.3 Correlation of Late Jurassic-Cretaceous Stage Boundaries to the GPTS

~0~

11.6 and 11.7). All the nannofloral datums have been directly correlated to polarity chrons. Ammonite biozones have been directly correlated to magnetic stratigraphy in parts of the Oxfordian-lowermost Valanginian (see Ogg et al., 1991c), Valanginian-Hauterivian (Channell et aL, 1995b), and uppermost Hauterivian-Barremian intervals (Cecca et al., 1994). The Oxfordian/Kimmeridgian boundary is correlative to the base of the Sutneria platynota ammonite zone, which has been correlated to CM25 in southern Spain (Ogg et al., 1984). The Kimmeridgian/Tithonian boundary is correlative to the base of the Hybonoticeras hybonotum ammonite zone, also in southern Spain (Ogg et aL, 1984). The Tithonian/Berriasian (Jurassic/Cretaceous) boundary does not have a universally accepted definition. Many of the candidate biomarkers have been correlated to the polarity chrons (Ogg and Lowrie, 1986; Channell and Grandesso, 1987; Bralower et al., 1989; Ogg et al., 1991c). The base of CM18 has become the generally accepted correlation to the Tithonian/ Berriasian boundary. The Berriasian/Valanginian boundary is defined by the base of the T. otopeda ammonite zone and falls within CM15n (Ogg et al., 1988) and between the first occurrence (FO) of Cretarhabdus angustiforatus and the FO of Calcicalathina oblongata (Fig. 11.6). The Valanginian/Hauterivian boundary coincides with the base of the A. radiatus ammonite zone and is close to the FO of Nannoconus bucheri and at the base of CM11n (Channell et al., 1995b). The base of the S. hugii ammonite zone defines the Hauterivian/Barremian boundary, which falls between the last occurrence (LO) of Lihtraphidites bollii and the LO of Calcicalathina oblongata and in the upper part of CM4 (Cecca et al., 1994; Channell et aL, 1995b). The Barremian/Aptian boundary was formally defined at the first occurrence of Deshayesites. None of the nannofossil events proposed by Thierstein (1973) to define this boundary have proved to be reliable. The FO of Rucinolithus irregularis is correlated to the upper part of the C. sarasini ammonite zone and is therefore slightly older than the Barremian/ Aptian boundary (Channell and Erba, 1992). The base of CM0 coincides closely to this boundary, but direct correlation of this polarity chron with the ammonite biozone has not been documented. The Santonian-Campanian is formally defined by the appearance of the ammonite Placenticeras bidorsatum, the index species of the oldest Campanian ammonite zone. Unfortunately, this species is extremely rare even in the type area (northwest Europe) and the Santonian-Campanian boundary is often informally defined on the basis of the FO of the nannofossil Broinsonia parca and/or the FOs of foraminifera Globotruncana arca and Bolivinoides strigillatus. The microfossil markers of the SantonianCampanian boundary result in a correlation of this boundary to the basal

204

11 Jurassic and Cretaceous Magnetic Stratigraphy

part of C33r, close to the top of the Cretaceous normal superchron (e.g., Alvarez et al., 1977; Channell et al., 1979; Monechi and Thierstein, 1985). The type area of the Campanian-Maastrictian boundary (northwest Europe) is characterized by a major hiatus, and hence there is more than average debate over the definition of this stage boundary. In the pelagic realm, the LO of the foraminifera Globotruncanita calcarata is often used as an informal definition, and this event usually appears in the top part of C33r (e.g., Alvarez et al., 1977).

Triassic and Paleozoic Magnetic Stratigraphy

12.1 Introduction The absence of preserved marine magnetic anomalies for times preceding the breakup of Pangea forces a change of strategy in the study of pre-Late Jurassic magnetic polarity stratigraphy. Irving and Pullaiah (1976) suggested the use of type sections to build the pre-Jurassic GPTS in the absence of the oceanic magnetic anomaly template. Up to now, type sections for magnetic stratigraphy have not been established; however, we believe that the type section approach would be helpful for pre-Late Jurassic time. Type sections for magnetic polarity sequences will have the same problems presently associated with type sections in classical stratigraphy, such as unconformities, poor age control, lack of diagnostic fossils, and inadequate descriptions. The ideal magnetostratigraphic type section would have not only a well-defined magnetic stratigraphy but also high-resolution biostratigraphy facilitating correlation to geologic stages. Many limestones with good biostratigraphic control have very weak magnetization intensities or are remagnetized. Red siltstones and sandstones, on the other hand, often carry a well-defined primary magnetization but are often difficult to correlate to the standard geologic stages due to lack of diagnostic fossils. In the Soviet Union, a hierarchical system of hyperzones (megachrons), superzones (superchrons), and zones (chrons) was established for Cambrian to Permian time (Khramov and Rodionov, 1981). Unfortunately, the original data on which this scheme was based are not available to Western researchers, and descriptions of type localities were not given. The proposed GPTS (Fig. 12.1) is, therefore, difficult to evaluate. In intervals for which 205

206

12 Triassic and Paleozoic Magnetic Stratigraphy

Modified

from KHRAMOV

Ma

1987

= . .:...

.:..

|

,,._

|

o,. oo ~ 250+

9| .,.,

,.J

z

~

9 .

,,.._

"~

=.

"~

|

' ~-

m

,

.

_< ,---

uJ

300

ca w

l~

~

:

C ~)

i

c

Z

"K

C)

--

_m_

<

r~*',

5

0 >

j !

N

|

";

l 1= l l

t

400"

|

~ tr -"

"~

=

=

>, -=

=

_

~.,:

Z

~

]

_.=

m m I::3 w

c

z

g,

~

"~

.

.=_ m e-

e

"9J i ~ 450"

1

" ~

.Z

I

I

9

~

rr i11 tt

1

l

'0

350"

E

r .u

l

_e

_<++._

-01~

1

Z

llil m

+

1

++ 1

-~

_,.

nn

._

m

500-

u.i

zl

550-

,..+

<

_.=

~

+

,<

~

|

"

l

,~ 5 " r l

E m

-

m .

F i g u r e 12.1 Geomagnetic polarity time scale derived from the Paleozoic of the former USSR. Normal polarity (black), reverse polarity (white) (after Khramov and Rodionov, 1981).

12.2 Triassic

207

more recent data are available, such as the middle Carboniferous, many more polarity chrons are known to exist.

12.2 Triassic Triassic magnetic stratigraphy was pioneered by Picard (1964) and Burek (1964, 1970) and expanded in clastic sediments from the western United States by C. Helsley and his students in the early 1970s (e.g., Helsley, 1969; Helsley and Steiner, 1974). Upper Triassic magnetic stratigraphy is best known from the lacustrine sediments of the Newark Basin. The initial magnetostratigraphic results from the Newark Basin were from industry boreholes and outcrops (McIntosh et al., 1985). The upper part of the sequence above the Watchung Lava flows, as well as the flows themselves, are considered Early Jurassic (Hettangian) in age on the basis of pollen stratigraphy. These units are normally magnetized and were labeled the Upper Normal Interval (McIntosh et al., 1985). This normal polarity zone extends down into the Upper Triassic (Rhaetian). Lower in the section the sediments were found to be characterized by mixed polarity over a stratigraphic thickness exceeding 2500m. Later studies by Witte and Kent (1989) and Witte et al. (1991), based on more complete demagnetization procedures and denser sampling, along with field tests for the age of the magnetization, led to a more complete documentation of the polarity sequence for sediments from the Newark Basin. From the Late Triassic Carnian stage to the earliest Jurassic Hettangian stage, at least six normal magnetozones were identified in the Newark Basin outcrops (Fig. 12.2, column 16). The excellent results obtained from the Newark Basin outcrops led to a study of drill cores. A series of seven drill sites were located to obtain a complete record of the sedimentary history of the basin. Core recovery was almost 100% and overlapping cores, each about a kilometer in length, resulted in a complete magnetostratigraphic record of lacustrine beds covering the time interval from the early Carnian to the Triassic/Jurassic boundary. The magnetostratigraphic study of these sediments is enhaced by the presence of a well-defined cyclostratigraphy tuned to Milankovitch frequencies. The stratigraphy was divided into lithologic units, called McLaughlin cycles, which are believed to represent the 413-ky eccentricity cycle of the Earth's orbit. An age of 201 Ma based on 4~ and U/Pb techniques has been obtained for the lowest flow unit of Watchung basalts at the top of the sedimentary sequence (Sutter, 1988; Dunning and Hodych, 1990). It is, therefore, possible to date events within the drill cores by counting down from the Watchung basalts using the cyclostratigraphy.

208

12 Triassic and Paleozoic Magnetic Stratigraphy

The magnetic stratigraphy of these cores has been studied by Kent et al. (1995). Samples were taken from the drill cores at 2-m intervals. All samples were analyzed by progressive thermal demagnetization and oriented using the direction of the magnetic overprint. The magnetic stratigraphy from the Weston core (Fig. 12.3) illustrates the high quality of the data. The sedimentary rocks in the Newark Basin dip to the north and the drill sites were located so that the stratigraphies from each core overlap. This has resulted in overlapping patterns of polarity zones (Fig. 12.4). The borehole studies confirmed the previously determined outcrop studies, but the increased resolution led to an increase in the number of polarity zones from 12 to 42, an increase of almost a factor of 4. The Milankovitch cyclostratigraphy places the sediment sequence in a precise time frame and yields a reversal frequency of 2 My -1 over late Carnian and Norian time (Fig. 12.2, column 19). Witte et al. (1991) have suggested that the Newark Basin qualifies as the type locality for magnetic polarity stratigraphy of the late Carnian and Norian interval. The lack of a marine biostratigraphy is offset by the essentially complete magnetic stratigraphy and the precise cyclostratigraphic timeframe. Middle and Late Triassic magnetic stratigraphies have been obtained from condensed marine "Hallstatt-type" reddish nodular limestones at two localities in southwest Turkey and from Austria (Gallet et al., 1992, 1993, 1994) (Fig. 12.2, columns 6, 9, 10, 17). The sedimentary sections are not thick ( - 3 0 m), and unconformities are present. The quality of the magnetostratigraphic data is excellent and field and laboratory tests for stability have been carried out. There are, however, discrepancies in correlation between the two Turkish sections which the authors attribute to the sections originating in different hemispheres. The southernmost section has been correlated to a third section in Austria, the polarity of which is thought to be unambiguous as it must have originated in the northern hemisphere. The polarity sequence is not easily correlated to that of the Newark Supergroup. Kent et al. (1995) have offered a partial correlation of the Turkish sections to the Newark Basin record. They correlate the Kavur Tepe magnetozones a+, e+, and g2+ (Fig. 12.2, column 9) to magnetozones d+, h+, and j+ (Fig. 12.2 columns 16, 19). They have also suggested a correlation of the Chinle section (Reeve and Helsley, 1972) to the Newark sequence. The Chinle section has four normal magnetozones (Fig. 12.2, column 1), the youngest of which (N4) is correlated to 1+ of the Newark sequence (Fig. 12, column 16). N2 and N3 of the Chinle section could be correlated to j + of the Newark sequence. The Carnian/Norian boundary based on pollen in the Newark sequence occurs between the upper part of polarity zone b+ and the base of polarity zone d+. This boundary, based on conodonts, is present in the sequence at

209

12.2 Triassic

TRIASSIC Ma

Stage

9

1

16

204

k-

208-

212-

l+

,~

!

ss S

I,,"

,"

e

"m )~ i-

UR,2

R r

h2

~/SIR11

gf+ e-

216.

iN11

a.

I

----2

I

IR10

220-

E13

N8

E12 Ell __

E:

A

,-

224-

~-

El0 E9 E8

,3

O

S 228" 232-

236-

S

I

B.~

u--.m

E7 E6

A(~:~ C"~ B A~:G~-

c~ E

I.U "~ _J _J

~__

~

~ 4

7

11

AB:

f-

C

<

240-

14 c ~-

~

rr

O I~ -o"

__J --C~

<

248"

E+

Dc+

I_ ~ . _~,._~

244-

20

,,,

(~' ~c

18

SpR1 SmN3 ~ N 6 N7RI R, ~ISmR2 ~SrnN2 r a N 5 15 ~SmR1 L._J R4, ~N3T SmN1 i~. I I ~,D,~R 2 ~ DDNz .R!~ - - r~~NI R 33 .,.. R ~ L__JGR2 I I R2 mN2T ~111~GN2 ~ "" " 'm~' N R~1' 'TT . . . iG I IGR1N,_~-. ~IN~N2R17 ..~

O

N

Figure 1 2 . 2 Magnetic stratigraphy of Triassic sections. Numbers refer to individual studies in Table 12.1. Right-hand column is composite section.

Bolucektasi Tepe just prior to a boundary from reverse to normal polarity, informally designated N i l (Fig. 12.2, column 6); prior to N8 in this column the quality of the magnetic stratigraphy deteriorates and the VGPs fall to low latitudes. Kent et aL (1995) tentatively correlated the base of polarity zone N i l to the base of b+ in the Newark magnetic stratigraphy; however, a secure pattern match is not possible. Unfortunately, the base of the

210 Table 12.1 Triassic Rock unit 1 Chinli

Lo Age Hi Norian

Region USA. NM

A

Q

+35.03 -104.08

NSE

NSI

NSA

M

2

.5-.9 m

1

145

2 Chinli

Carnian

USA, N M

-106.5

3

240 20

3 Fleming Fjord

Carnian

1

134

1

12.2

4 Moenkopi

Anisian

East +71.73 -23.41 Greenland +35 -106.5 USA, NM

4

75

5

110

+35

5

55

D

RM

A

NMZ NCh %R

RT

F.C.T

Q

DM

AD

References

- 1-0

T

B

F

9

5

49

R+

1-0

T

Z-P F

5

3

51

R-

T

B

5

-

T

Z-P F

19

6

51

R-

- -

-

B

15

12

62

-

- I-TdI

A-T Z-P F-V

23

-

26

R+

8 Molino-Garza er a/. (1991) 5 Helsey and Steiner (1974) 7 Gallet e r n / . (1992)

5

3

84

R+

6 Purucker er a/.

23 15 8 6

5 -

80 50 40 46

R+

-

Shive et 01. (1984) 8 Gallet e r a / . (1993) 8 Gallet er a/. (1994) 6 Steiner and Lucas

Reeve and Helsley (1972) 7 Moho-Garza er a/. 5

(1991)

5

Moenkopi

F. Triassic

+38.58

-108.93

3

.35 m

1

150

NorianCarnian Ani. Spa.

Turkey

+37

+30

1

262

1

72

USA. A n

+35.8

-111.45

2

1

118

8 Chugwater Fm. 9 Kavur Tepe 10 Mayerling 11 Anton Chico

E. Triassic Norian Carnian Anisian

USA Turkey Austria USA, NM

+43 -107.00 +38 +30.5 +48.05 +16.13 +34.9 +105.7

5 1

2 2

.3 m 800 563 179 67 87

12 Feixianguan Fm. 13 Mezarlik

Die-Greis.

China

+32

1

Carnian

Turkey

+36.53 +30.56

1

6 Boliicektasi Tepe 7 Moenkopi

- ~

1-0

F-D-I

F-D-I

1

F D-T F-D-I 179 - J T B 30 - I-TdI T-A 2 - P F-V 25 I T 2 - P F-D-I 23 - 0 T Z-P F-D-I

200

1

130

53

1

5

1 1 1

- -

T

B

~

6 0 -

3 Reeve el a/. (1974)

(1980)

~

-

R+

7

(1992)

-105.5

~

-

1

T

2 - P F-V

1

T

Z-P F-D-I

6

-

30

R-

7

14

-

47

R+

8 Gallet e r a / . (1994)

Heller er a/. (1988)

14 Chios 15 Russia I6 17 18

1Y

20 21

Newark Supergroup Bolucektasi Tepe ArcticArchipelago Newark Supergroup Hydra Feixianguan Fm. Wombat Plateau Iberia

Anisian

Greeve Russia

+37.31

t22.82

150

USA

+40.25

-75.25

24

HetCarnian Norian

Turkey

Spt.-Grie

Canada

Het: Norian AnisianLadinian Spat-Gries

USA

+40.75

-75.5

Greece

+37.3

t23.43

China

+2Y.Y

+ IM.3

Australia

- 16.95

+ I15.6l

NormCarnian ZeohCarnian

13 1600

I25

T

Z-P F-V F-DI

3 7

T

2-P F-V

4

T

Z-P F-V

T

-

4180 K

I

R+ R-

Ft

Y

57

-

3

F+

Y

55

R+

23

-

26

R+

2 - P F-DI

27

17

62

R+

F+

Y

T

Z-P F-I

45

12

51

R+

F+

II)

T

2-P F-V

X

52

Z-P F-DI

+ 37

+30

262

I

72

+80.5

-95

342

1

I100

1695

1

4660 K

36

1

23

4

501

1

Y60

T

I6

2

356

400

A-T Z-P I

I6

4

200

Yo0

T-A 2-P F

10

37

-~

I

Muttoni el al (1995) Gurevich and Slautsitays (lY85) Witte PI a/. (1991) Gallet

?I

a/. (1YY2)

Ogg and Steiner (1991) Kent er o l (1995)

-

Spain

-

I

R+(c)

-

7

65

R+

F+

9 Steiner er a/. (IYXY)

63

-

Muttoni er al (IW)

6 Galbmn cr d.(1W2) -

4 Turner rral. (19x9)

211

|2

12 Triassic and Paleozoic Magnetic Stratigraphy

Stratigraphy Chrono-

VGP Latitude

Litho-90 ~ OO

0o

Polarity 90 ~ -0 Meters E21n

NN

100 9

MM

- 200

E20r

LL KK O

- 300

~5

r

#

(9

I

JJ

-400

I

II

E20n E19r

-5OO

Ukrainian Cedar Grove

)

I

-600 )

FF

/

-700

E19n E18r

E18n

i-

EE DD

E17r 800

Figure 12.3

Magneticstratigraphy of the Weston core, Newark Basin (after Kent et al., 1995). Polarity chrons are labeled as normal and reverse pairs, with the reverse overlying normal.

Carnian is not identified in the Newark basin. In the Mayerling section (Austria) (Fig. 12.2, column 10), the Ladinian/Carnian boundary occurs in the normal polarity zone A + . At other sections, this boundary occurs at unconformities (Gallet et al., 1994). Middle Triassic magnetic stratigraphy is less well established than that of the Late Triassic. In the Middle Triassic Moenkopi Fm. of Arizona and New Mexico, the biostratigraphy must rely largely on rare vertebrate finds. The magnetic stratigraphy of the Moenkopi Formation is well known

Figure 12.4 Correlation of Triassic cores from the Newark Basin (after Kent el al., 1995). Inset indicates location of cores in Newark Basin. VGP latitudes are depicted by a line passing through the data points.

|4

12 Triassic and Paleozoic Magnetic Stratigraphy

through the studies of Helsley (1969), Helsley and Steiner (1974), Purucker et al. (1980), and Molina-Garza et al. (1991). Steiner et al. (1993) place the

upper part of this formation in Arizona and Colorado in the early Anisian. Molina-Garza et al. (1991), on the other hand, place a section of the Moenkopi from New Mexico in the late Anisian/Ladinian, based on fossil vertebrates (Fig. 12.2, column 4). Steiner and Lucas (1992) have obtained a magnetic stratigraphy from the Anton Chico member of the Moenkopi Fm. which they place in the Anisian on the basis of vertebrate fossils (Fig. 12.2, column 11). The polarity zone patterns from the Anton Chico member given by the two studies from New Mexico are very similar and are said to correlate to the Holbrook member of the Moenkopi Fm. in Arizona. The only studies of marine sections from the Middle Triassic are by Muttoni et al. (1994) and Muttoni and Kent (1994). The first of these studies delineates the magnetic stratigraphy across the Anisian/Ladinian boundary from a section on Hydra Island (Greece). The quality of the data is good and the stage boundary occurs within a normal polarity zone informally designated C+ (Fig. 12.2. column 20). The other study is on limestones from the Southern Alps and the section is in sediments of the latest Anisian age. The section studied is entirely normally magnetized and probably correlates with magnetozone C+ of the Hydra section. Magnetostratigraphic results have been reported from the Middle Triassic Muschelkalk Fm. of Spain by Turner et al. (1989). The quality of the magnetostratigraphic data from this study is poor. A large proportion (85%) of the samples did not yield interpretable demagnetization data, and this magnetic stratigraphy is not included in Figure 12.2. Lower Triassic results have been obtained by Steiner et al. (1989) from the Feixianguan and Jialingiang formations of Sichuan (China) from a 1-kmthick section which ranges in age from the Permo-Triassic boundary to the Spathian (upper part of Lower Triassic) (Fig. 12.2, column 21). In a recent compilation (Steiner et al., 1993), stratigraphic gaps are shown in the Spathian, but how these are recognized is not clear. Ogg and Steiner (1991) have studied the stratotype sections of the Early Triassic from Ellesmere and Axel Heiberg islands in the Canadian Arctic. The data quality of the three sections is not as high as that from the Sichuan; however, the magnetic stratigraphy is readily interpretable (Fig. 12.2, column 18). If these two sections cover the entire Early Triassic then the correlation of polarity zones would seem to be straightforward, since nine normal polarity zones are present in each section (Fig. 12.2). More detailed biostratigraphy is needed to solve the problem of possible stratigraphic gaps in the late Early Triassic of Sichuan. Unfortunately the boreal fauna ofthe Arctic cannot be easily correlated to the tropical fauna of the Chinese sections, inhibiting unequivocal magnetostratigraphic correlation. The magnetic stratigraphy of the boundary of the Early/Middle

12.3 Permian

215

Triassic (Spathian/Anisian) has been studied by Muttoni et al. (1995) in fossiliferous limestones from Chios Island (Greece). The stage boundary, based on ammonites, falls within normal magnetozone C+ (Fig. 12.2, column 14). The authors correlate normal magnetozone C + at Chios with normal polarity zone SpN2 in the Arctic sections and with N8 in the Sichuan section, although a correlation to N9 appears more likely. Magnetostratigraphic studies across the Permo-Triassic boundary indicate that a polarity zone boundary from reverse to normal is present at, or shortly above, the Permo-Triassic boundary. Studies in the former Soviet Union and China (Steiner et al., 1989; Heller et al., 1988) indicate that the basal Triassic normal polarity zone corresponds to most of early Griesbachian time. McElhinny (1973), in an analysis of polarity bias, concluded on the basis of data available at that time that the geomagnetic field during Triassic time was 75% normal polarity. In Figure 12.2, the right-hand column represents a composite magnetic stratigraphy, and it can be seen that the magnetic field is not biased toward normal polarity. The reversal sequence given here for the Triassic is most reliable for the Early Triassic, upper Carnian, and Norian times and least secure in the Middle Triassic.

12.3 Permian Irving and Parry (1963) were the first to recognize the long period of reverse polarity spanning Late Carboniferous and Permian time. They named it the Kiaman Magnetic Interval, after the name of the town in New South Wales (Australia) where rocks showing reverse polarity were first reported by Mercanton (1926). The name Kiaman has been retained by Russian workers but was abandoned by Irving and Pullaiah (1976) and replaced by the PermoCarboniferous Reverse Superchron (PCRS). The PCRS is known to end within the Late Permian. Irving and Parry (1963) originally placed the termination of the Kiaman at the beginning of the Narabeen Chocolate Shales of the Sidney Basin (Australia), which are Late Permian and Early Triassic in age. They called this polarity change the Illawara reversal. A restudy of this formation was carried out by Embleton and McDonnell (1981), who confirmed a reverse to normal polarity change in red claystones near the base of the Triassic. The older sediments in this sequence are unstably magnetized so a complete magnetic stratigraphy could not be established. The Upper Permian and Lower Triassic sequences in North America, Europe, and in many of the Gondwanaland continents often contain important unconformities. Upper Permian marine sediments are present, however, along the northern margin of Gondwanaland (Pakistan and India) as well as

216

12 Triassic and Paleozoic Magnetic Stratigraphy

in China. The best quality magnetic stratigraphy for later Permian time comes from the Wargal and Chidru Formations of the Salt Range in northern Pakistan (Haag and Heller, 1991) (Fig. 12.5, column 13). These sections comprise mainly carbonate sediments and yield an excellent fauna allowing correlation to other sections. The magnetostratigraphic record spans late Kazanian and Tartarian time, with six normal and five reverse polarity zones. The top of the PCRS was not determined as the lower part of the Wargal Formation is normally magnetized. Unfortunately, it has become apparent that there is an unconformity at the top of the Permian in this section; sediments equivalent to Changxingian time in the Chinese sequence are apparently missing (Wignail and Hallam, 1993). The biostratigraphic and magnetostratigraphic correlation of the Salt Range sequences to China (Fig. 12.5, columns 11 and 12) (Steiner et al., 1989; Heller et al., 1988) is therefore tenuous, although the correlation of the thick normal polarity zone near the top of the sequences in both China and Pakistan is reasonable (Fig. 12.5).

Ma

I

PERMIAN USA

RUSSIA

CHINA

~

2 6 0 ~ .-

~

m

3

13

m

270-gc i.

c

g

290-"~, E

"K 7 .

6

.

.

.

8 .

ui

3OO

m

m

Figure 12.5 Magnetic stratigraphy of Permian sections. Numbers refer to individual studies in Table 12.2. Right-hand column is composite section.

Table 12.2 Permian Rock unit

Lo Age Hi

Region

0

A

NSE

NSI

NSA

M

D

-

-

1 Tartarian

Tartarian

E. USSR

2 Tartarian

L Tri-Tai

USSR

3 Tartarian 4 Dewey Lake

Tartarian-Kz Ochoan

-.

-

+34.5

-101.5

3

19

5

52

5 Cutler Fm.

Permian

Urals USA. W. Tex. USA. Utah

+38.41

- 109.58

1

.6 m

1

2

2.5-5 m 760

I

1500

-

-

-

-

- V

57

R-

5

-

-

4

.34

R+

200

- J

T

B

1

100

5

- -

T

Z-P

F DI F-D

I

100

6

1-5

700

7

Casper Fm.

USA. Wyo

t41.45

-105.21

1

190

lngleside Fm

USA. Wyo.

+40.8

-105.2

1

37 m 549 243

I

8

E. PermianL. Carb. E. PermianL. Carb. Acch-Rot.

1

70

Germany

+51

+ 10.30

-

3

4000

Tri~ Permian Lopingian

China, Tarim China, Sich. China, Sich. Pakistan

+42.13

-83.35

1

21

5

870

+32

+I052

1

200

1

130 S

1-32.4

-106.4

2

85

-32.6

+71.n

1

220

Wargal and Chidru

- Molostovsky

-

68

13

-

F

-

References

-

2-P

1

ChangsingWujiaping Zechstein

0

-

106.4

12 Wujiaping

F.C.T

-

-

Dalong Fm.

6

%R R T

T

+39.73

11

F

NCh

-

USA, Colo

Tarim

B

NMZ

I

E. Permian

10

T

A

K

Maroon-Minturn Fm.

40

AD

__

h

9 Rotleigande

RM DM

-

9

5 - J-I

T

2-B

F

1

-

5

-

T-A

V

F

7

95 -

4

- -

T

Z-P-G

F

3

R+

5

I

T

Z-P

F-V

6

1

120 - I

T

2-P

F-V

7

-

I

2ol) --

1

T

Z-P

F-V

10

4

.72

29 R -

6

.5R

R-

7

38

-

6

(1992) Gurevich and Slautsitays (1985) Khramov (1974) Molina-Garza el a/. (1989) G n e and Helsley (1972) Millei and Opdyke (1985) Diehl and Shive (1981) Diehl and Shive (1979) Menning ec a/. ( 1988) McFadden era/. ( I 988) Heller el a/. (1988) Steiner er a/. (1989) Haag and Heller (1991)

| 13

12 Triassic and Paleozoic Magnetic Stratigraphy

The correlation of the Upper Permian marine sequence in Pakistan to nonmarine red beds which dominate the rock record in Russia, Western Europe, and North America and South America is not clear. A recent study of the Ochoan (Tartarian) Dewey Lake Formation in northern Texas (Molina-Garza et al., 1989) shows at least two normal polarity zones in the Late Permian (Fig. 12.5 column 4). Khramov (1974) shows at least five normal polarity zones in sediments from the Upper Permian type section of the Ural Mountains (Fig. 12.5, column 3). The data quality appears to be high, and the results are internally consistent. Unfortunately, these sections have an unconformity in uppermost Permian and lowermost Triassic, similar to the situation in the western United States. A considerable amount of magnetostratigraphic work has been carried out on the Permian Rotliegende and Zechstein Formations of Germany, both in outcrop and bore holes (Menning et al., 1988) (Fig. 12.5, column 9). There are normal polarity zones in the upper Rotliegende and throughout the Zechstein Formation. Menning et al. (1988) showed at least seven normal polarity zones in the Late Permian. This number of normal polarity zones is more or less consistent with the results from the former USSR and Pakistan, where five and six normal polarity zones have been recorded in this interval, respectively. In the composite polarity column for the Permian stage (Fig. 12.5, right column), eight normal magnetozones are given for Late Permian time, derived principally from the records from China, Pakistan, and central Europe. The position of the top of the PCRS is older than the Tartarian of Russia, the Zechstein and upper Rotliegende Formations of central Europe, and the base of the Wargal Formation in the Salt Range of Pakistan. The Wargal Formation is Murgabian in age, correlative to the Guadalupian or Wordian of the North America and to the base of the Kazanian stage in Russia (Haag and Heller, 1991; Catalano et al., 1991). The upper boundary of the PCRS in North America lies beneath Ochoan units of west Texas, which have an associated 4~ age of 251 _ 4 Ma on volcanic ash (Molina-Garza et al., 1989). The boundary between the Permian and Triassic in China has been dated at 251 Ma by Claou6-Long et al. (1991) using the SHRIMP ion probe at the Australia National University. Taken at face value, the date from Texas would place the Ochoan magnetic stratigraphy very close to the Permian-Triassic boundary. Harland et al. (1990) have subdivided the Permian stage into the Rotliegende and Zechstein epochs in a bipartite division. The base of the Zechstein Epoch is given a numerical age of 256 Ma by Harland et al. (1990). Valencio et al. (1977) in Argentina have reported a normal polarity zone lying above an intrusion dated at 2 5 9 _ 7 Ma which they called the Q. Del Pimiento event. In the same region, Creer et al. (1971) also noted normally magnetized lavas at 258 • 5

12.3 Permian

219

Ma. It is possible that these normal polarity zones may mark the end of the PCRS and are not subchrons within the PCRS. Further sampling in the Argentine sections may offer an opportunity to date the end of the PCRS. Due to the lack of reliable radiometric ages directly associated with sediments of Late Permian age, stage boundaries of the Late Permian are poorly constrained in time. The precise date of the end of the PCRS is therefore unknown. Haag and Heller (1991) believed it to be older than 261 Ma, whereas M. Menning (personal communication, 1992) gives an age of 265 Ma. In Figure 12.5, an age of 262 Ma is given for this boundary; however, it should be understood that nowhere is it directly dated. Over the years, reports of short normal subchrons within the PCRS have been published in both Russian and Western literature. Helsley (1965) found normally magnetized sediments in the Dunkard Formation of West Virginia, which is believed to be Early Permian in age. The initial study was done with alternating field demagnetization, which is notoriously ineffective for removal of secondary overprints from hematite-bearing samples. Samples from the normal polarity horizon were subsequently subjected to thermal demagnetization by Gose and Helsley (1972) and were found to retain their normal polarity. Normal polarity magnetizations have been obtained from Lower Permian Pictou sediments of Prince Edward Island by Symonds (1990). The normal polarity zone in the Pictou beds may be correlative to that in the Dunkard Fm. of West Virginia. Menning et al. (1988), in their study of the Rotliegende Formation of eastern Germany, found normally magnetized sediments in the Early Permian part of this sequence which they placed at the Permian/Pennsylvanian boundary, implying a correlation with the Dunkard series of West Virginia. It does seem certain that one or more normal subchrons occur within lower Permian strata, although it is unclear whether the normal polarity zones recognized at different locations are correlative. Several sections have been sampled across the Permo-Carboniferous boundary in the western United States (Diehl and Shive, 1979, 1981; Miller and Opdyke, 1985) and no normally magnetized rocks were detected. Sinito et al. (1979) gave evidence for a short subchron within the PCRS, positioned 100 m below a lava dated at 288 _+ 5 Ma. This may be equivalent to the normal polarity subchron seen in the Dunkard Fro., although it was originally thought to be the base of the PCRS. Normally magnetized sites of Early Permian age have been reported in several tectonic studies (Halvorson et al., 1989; Wynne et al., 1983; Irving and Monger, 1987) and may be correlative with the Dunkard subchron. Although normal subchrons may occur within the PCRS, magnetostratigraphic studies by Diehl and Shive (1979, 1981), Miller and Opdyke (1985), and Magnus and Opdyke (1991)

220

12 Triassic and Paleozoic Magnetic Stratigraphy

have shown that the PCRS is dominantly reverse polarity, as claimed by Irving and Pullaiah (1976) (Fig. 12.5, columns 6-9).

12.4 Carboniferous The age of the base of the PCRS has not been resolved. In North America, McMahon and Strangway (1968) reported normal polarity zones in sediments of the Upper Demoinesian Minturn Formation of Colorado. The section was restudied by Miller and Opdyke (1985), who did not find any normal polarity zones in the section. The study was extended throughout the Minturn Formation by Magnus and Opdyke (1991), who determined that the reverse polarity zone (PCRS) extended to the base of the Minturn Fm., which is earliest Demoinesian or late Atokan in age. Roy and Morris (1983) had shown that sediments with dual polarity are present in rocks as young as Westphalian A in the Maritime provinces of Canada. In the same region, DiVenere and Opdyke (1990, 1991b) documented an internally coherent set of polarity zones in the Maringouin, Shepody, and Claremont Formations of Namurian and Westphalian A age. At least seven normal polarity zones are present, giving a reversal rate of about 1 My -~ (Fig. 12.6). The age of magnetization is constrained to the Late Carboniferous by a positive fold test. An attempt was made to extend the magnetic stratigraphy to younger formations exposed along the shores of the Bay of Fundy. All samples were reversely magnetized; however, a data gap of 700 m occurs in the sequence (Fig. 12.6). These data yield the best and most complete reversal sequence now available for the early Late Carboniferous. The dual polarity for rocks of early Pennsylvanian age is supported by a study on red Morrowan paleosols from Arizona (Nick et aL, 1991). The age of the beginning of the PCRS was originally determined to be 310 Ma based on the K/Ar age of the Paterson Volcanics in Australia. These volcanics were shown by Irving (1966) to be normally magnetized and were used to define as the base of the PCRS (Kiaman). Sites taken from the Seaham Formation, directly overlying the Paterson Volcanics, were reversely magnetized and were thought to be Late Carboniferous in age. Upper Carboniferous sites from the Currabubula and Lark Hill Formations were also reversely magnetized and passed the fold test. The base of the Kiaman was, therefore, thought to be firmly determined. However, a recent study of the Paterson Volcanics using single-crystal zircon dating by Claou6-Long et al. (1995) has yielded a date of 328 Ma, which places them within the Visean, much older than the normally magnetized sequences from the Canadian Maritimes (Fig. 12.6). Therefore, the Paterson

221

12.4 Carboniferous

Ma

3000 m

t. <

Canadian Maritimes 2000 m

Nova Scotia

New Brunswick

Mauch Chunk Formation Pennsylvania

--....

-90 1000 m

Pottsville

9O

Lavelle

Jim Thorpe

(DiVenere and Opdyke, 1990)

-90

90

(DiVenere and Opdyke, 1991a)

-90

90

(DiVenere and Opdyke, 1991 b)

S

I

I

I

I

I

I

N

Figure 12.6 Magnetic stratigraphy of the Canadian Maritime provinces and the Mauch Chunk Fm. (Pennsylvania).

222

12 Triassic and Paleozoic Magnetic Stratigraphy

Volcanics cannot be used to fix the onset of the PCRS. Recent studies in .Australia (Opdyke et al., 1960) have located the base of the Kiaman and dated it at approximately 320 Ma (Claou6-Long, personal communication). The other significant data set that bears on this problem is the magnetic stratigraphy of the Middle Carboniferous strata of the Donetz Basin in the Ukraine. Khramov (1974) has presented a magnetic stratigraphy for the Upper Carboniferous which shows normal polarity zones as young as Upper Moscovian. Opdyke et al. (1993) restudied these sediments, and although it was possible to recover a prefolding magnetization which is probably Carboniferous in age, it was not possible to confirm Khramov's original magnetic stratigraphy. At this time, the only unimpeachable data that bear on the age of the base of the PCRS are those from Maritime Canada. The highest normal polarity zone in this sequence occurs within sediments of Namurian age and yields a duration for the PCRS of approximately 58 My (from 262 to 320 Ma). In North America, the reversal pattern determined in the Lower Pennsylvanian has been extended into the Upper Mississippian (Visean) in sediments of the Mauch Chunk Formation of Pennsylvania (DiVenere and Opdyke, 1991a) (Fig. 12.6). Opdyke and DiVenere (1994) have reported a minimum of nine normal polarity zones within the Mauch Chunk Formation (Fig. 12.7, column 7) with a reversal frequency of 1.5 My -1, similar to that of the early Pennsylvanian. The geomagnetic field appears to have been about 50% normal and 50% reverse in Middle Carboniferous time (Fig. 12.7). Little is known of the polarity pattern for the rest of Early Carboniferous time; however, both polarities are present in sediments of Visean and Tournaisian age (Irving and Strong, 1985). The beginnings of a magnetic stratigraphy has been obtained from Lower Carboniferous lavas of the Midland Valley of Scotland of Torsvik et al. (1989) (Fig. 12.7, column 8). Turner et al. (1979) and Palmer et al. (1985) have reported magnetic stratigraphy from Carboniferous limestones from Great Britain; however, these limestones have been remagnetized (McCabe and Channell, 1994) and the magnetic stratigraphy is therefore unreliable.

12.5 Pre-Carboniferous Middle and Early Paleozoic magnetic stratigraphies are largely restricted to the former Soviet Union. Lower Paleozoic data are mainly from the Siberian platform. In North America and Western Europe, severe remagnetization problems have resulted in very few reliable paleomagnetic pole positions and a paucity of magnetostratigraphic data. A very condensed sequence across the Famennian/Frasnian boundary of the Late Devonian appears to record several reversals of the geomagnetic

223

12.5 P r e - C a r b o n i f e r o u s

CARBONIFEROUS

My

m ~

N

Q,

5

"E .=31o- ~ c

i

~;

d

~

~

i

@

m

era O

m

O L._

z

ml

II

m m

m m

m III ~

~

i

6

II

II

II

II

m

m

7 m m

c

330"

.m_ =

m

E {o ,,.

m

m m

=

m

-

9

8

~. 9 > r,l

._~ 340- . m g

II

II

II

/I

c @

c

cg m

.m_ o t~ c

t,, cg

"3 c ,,._

3 o

~

I,-.

c

~= I-,-

Figure 12.7 Magnetic stratigraphy of Carboniferous sections. Right-hand column is a composite. N u m b e r s refer to individual studies in Table 12.3.

field (Hurley and Van der Voo, 1990). A reversal sequence of this age does not appear in the Russian GPTS. Douglas (1988) has studied an --~1 km thick section of the Ringerike Fm. of Norway which is considered to be Ludlow (Late Silurian) in age. Two formations were studied, the Sundvollen Fm. and Stubdal Fm. Most of the Sundvollen Fm. appears to be normally magnetized while the uppermost Sundvollen Fm. and overlying Stubdal

Table 12.3 Carboniferous Rock unit 1 Supai Gp.

Lo Age Hi

Region

NSE

A

NSI NSA

M

D

190 - J-I

T

2-V

I:

1

1

100 .-

700

T

2-P

F-D

1

I

100 --

1

2-P

F

1

1

100 --

-

~r

2-P F-v

9

-

-

T

2-P

II

-

+41/45

-105.21

1

1

+39.73

-106.4

II

.37 549 68

1-5

+38.2

-105.85

1

55

3

1280 -

+45.75

-64.38

I

177

I

2100

-

+45.75

-64.29

1

.5

1

600

-

+40.75

-76.42

Vir-Des

USA.

RT

100 --

I

Minturn Fm.

%R

1

145

3

NCh

I

3

E.P.-L.C

MZ

F

+35.2N -113

Casper Fm.

A

2-P

USA,

2

AD

T

Virgillian

Arizona USA, Wyoming

RM DM

48

-

~

~

-

F.C.T

Q

References

6

Steiner (1988)

5

Diehl and Shive (1981)

-

5

Miller and Opdyke (1985)

C+

6

Magnus and Opdyke (1991)

R + F+

7

DiVenere and Opdyke (1991)

7

Divrnere and Opdyke (1990)

~

Colorado 4 5

Minturn Fm. Cumberland Grp.

MissAtoka Na. West

Cumberland Grp.

Na. West

6

Mauch Chunk Fm.

Late Miss.

7

Mauch Chunk Fm.

Late Miss.

8

Kinghorn

E. Carb.

USA,

Colorado N. Sco.. Canada Canada, N.B. USA, Penn. USA. Penn. U.K., Scotland

+56.2

-3.1

~

I:-V

78

5 R+

F+

1

105

1

340 - -

T

2-P

F-V

6

-

41

R + F+

7

Divenere and Opdyke (1991a)

2

164

1

550

-

T

2-P

F-V

22

-

53

R+

F+C+

7

Opdyke and DiVenerc (1994)

1

16

5

566

K

T-A

2-P

F

46

K+

-

8

Torsvik

~

J-K

4

~

el

o l (1989)

12.5 Pre-Carboniferous

225

Fm. are reversely magnetized. The normal polarity zone is well documented and the high unblocking temperature magnetization is most probably Silurian in age. Trench et al. (1993) have compiled Silurian paleomagnetic data from all continents. They identified 19 studies which they considered yield reliable polarity information. The compilation of this data set yields mixed polarity for early and later Silurian time; however, rocks of Wenlock (middle Silurian) age yield only normal polarities. This study provides an indication of reversal pattern but this type of analysis cannot take the place of magnetostratigraphic study of individual sections. It should also be noted that this compilation is in serious conflict with the Russian GPTS (Fig. 12.1). For the Lower Paleozoic, the magnetic stratigraphy of the Ordovician is best known. Early work on Ordovician magnetic stratigraphy was reported by Rodionov (1966) and Khramov (1974) (Fig. 12.1). These early studies employed temperatures of 100~ and alternating peak fields of less than 10 mT. This low-level treatment is clearly inadequate by modern standards, and the reverse and normal directions are not 180~ apart, indicating unresolved magnetic overprints. Nevertheless, it appears that the directions record two polarities and that the observed polarity zones are reproducible on a regional scale across the Siberian platform. A second important study, but of more recent vintage, has been carried out on Ordovician (Llanvirn-Caradoc) limestones from Vastergotland, southern Sweden (Torsvik and Trench, 1991). The composite section is about 26.5 m in thickness and therefore very condensed. An internally consistent sequence of normal and reverse polarity zones has been obtained (Fig. 12.8, column 4). The top of the sequence, Caradoc in age, is normally magnetized. The Llandeilo is mainly reversely magnetized and is separated from the overlying and underlying parts of the sequence by gaps in data. The basal 9 m of section is Llanvirn in age and, although dominantly reversely magnetized, contains two normal polarity zones. Trench et al. (1991) have reviewed the polarity of Ordovician paleomagnetic data worldwide. They point out that a study of the Upper Ordovician (Ashgill) Juniata Fm. of Pennsylvania by Miller and Kent (1989) indicates entirely normal polarity. Fourteen other studies which were designed to produce data for pole positions, rather than for magnetostratigraphic purposes, were judged by Trench et al. (1991) to possess reliable polarity information. These studies help to ratify the magnetic stratigraphy from the primary sections (Fig. 12.8, columns 1,4). An important result of these correlations is a reassessment of the biostratigraphic correlation between the Baltic Shield and the Siberian platform (Chugaeva, 1976; Trench et aL, 1991). The Ordovician Period covers ---60 My and from data now available, it is certain that the magnetic polarity record for the Ordovician is much

226

12 Triassic and Paleozoic Magnetic Stratigraphy

ORDOVICIAN 440-

Ashgill

,

1

II II

450R

Caradoc

460- D

Llandeilo 470-

V

ml Llanvirn

I

480Arenig

490-

I

500- N Tremadoc

510

.r

......

D

Figure 12.8 Magnetic stratigraphy of Ordovician sections. Right-hand column is a composite. Numbers refer to individual studies in Table 12.4.

Table 12.4 Carnbro-Ordovician Rockunit I

Lena River

Lo Age 131

Region

A

Q,

NSE

NSI

CoradocTremedoc

Russia

Ashgill Lungwllm

IJSA. Penn. N. Wales

+40.5 i 53. I 7

-78 3 so

21 69

CaradocLlanvem Middle Ord. Low. Ord. LI-Arenig

Sweden

+SX.3

-139

43

NSA

M

D

RM DM

AD

A

MZ

NCh

9cR

RT

FCT

Q -

Reference\ Khramov er 01. (1965). cited in Trench Pr ol. (1991)

2 Juanita Fm. 3 Llanbttdrog and Mynytho G p s 4 Vustergotland

Lena River 6 Siberian Platform 7 Everton Fm. S

Black Mtn.

E.O.~L.C.

I

20

-

-

K

J 0

- J

A

Z-P V

F F

I 1

T

2-P

F-D

6

-

65

B DI 2-P F

24 16 I6

-

42 36 47

2-P F-D

13

-

67

-

F+ F+

5 6 7

I Russia Russia USA. Arkansas Queensland. Aus.

23 I 360

1 1 1-3

I69

I

I

-

-

+36.2

-92.66

122.6

- 14.03

I

39 90 114.5 I(XM

-

3 8

7

Miller and Kent (1989) Thomas and Bridm (1976) Torsvick and Trench (1991) Khramov (1974) Metallova er nl. (1984) Farr er nl. (1993) Ripperdan and Kirschvink (1992)

2~8

12 Triassic and Paleozoic Magnetic Stratigraphy

more complex than that proposed by Trench et al. (1991). A new study of Ordovician rocks from Arkansas (Farr et al., 1993) has yielded 16 polarity zones spanning late Arenig and Llanvirn time (Fig. 12.8, column 7) and a section presented by Khramov for the Middle Ordovician of the Lena River also shows many (---22) polarity zones. Unfortunately, faunal information is lacking from the Lena River section. What seems to be certain, however, is that the Middle Ordovician is a time of relatively frequent reversal of the geomagnetic field (Fig. 12.8). A study of a Lower Ordovician section from the Lena River yields at least 16 polarity zones in Tremodoc sediments (Metallova et al., 1984) (Fig. 12.9; Fig. 12.8, column 6). The base of the Early Ordovician appears to be reversely magnetized in this section. It is clear, when the original data are compared to the polarity scale presented by Khramov and Rodionov (1981) (Fig. 12.1), that their scale was greatly simplified. The most important magnetic reversal studies in rocks of Cambrian age are from the Siberian Platform (Khramov, 1974; Kirschvink and Rozanov, 1984). The stratigraphy presented by Khramov spans the time interval from the middle Cambrian into the Ordovician. Data are now available from Australia from an expanded section at Black Mountain (Queensland) for Late Cambrian and Early Ordovician time (Ripperdan and Kirschvink, 1992) (Fig. 12.10). The section in Australia is 1 km thick, while the Siberian sections rarely exceed 100 m in thickness. An unambiguous correlation between these two regions is not possible, but there are some common features in the magnetostratigraphic records. Both studies show reverse polarity at the Cambro-Ordovician boundary and both show five normal polarity zones in the Late Cambrian. A more definitive correlation between the two regions will require further study of the Russian sequences and a more complete presentation of the biostratigraphy of the Siberian sections. Kirschvink and Rozanov (1984) have analyzed sediments from the lowermost Cambrian of the Siberian Platform, and a magnetic stratigraphy across the TommotianAtdabanian boundary was obtained. The magnetization components pass a fold test; however, they yield a pole position which is far away from the Siberian apparent polar wander path, making polarity designations ambiguous. Kirschvink et al. (1991) have inverted the initial polarity designations based on changing ideas of paleogeography. The correlation of this pattern to a similar age record from the Amadeus Basin of Australia (Kirschvink, 1978) is not possible, perhaps because of unconformities in the Australian record. A magnetostratigraphic record from this time interval given by Wu et al. (1989) is now thought by the authors to be the result of remagnetization.

12.6 Polarity Bias in the Phanerozoic

RIVER LENA EARLY M

10

150

210

229

SIBERIA

ORDOVICIAN D~

270

o

330

i -30 0 3060

Figure 12.9 Magnetic stratigraphic study of Early Ordovician sediments from the Lena River, Siberia (after Metallova et al., 1984). D and I denote declination and inclination, respectively. Polarity interpretation is reproduced in Fig. 12.8, column 6.

12.6 Polarity Bias in the Phanerozoic The reversal record of the geomagnetic field from the Carboniferous to present is known well enough to allow us to examine the overall reversal pattern for the last 330 My of Earth history. It was previously thought that long segments of this record were characterized by polarity bias. Except

230

12 Triassic and Paleozoic Magnetic Stratigraphy

DECLINATION Datson Mbr

1000

Co e Member Z

t

~

!

0

Mort Member

i

~

! !

Jl I

angulatus

Cordylodus lindstromi

800--

"

700--

Jiggamore Member

:

:

i

i 500--

[]

I

i

600--~ . . . . . Unbunmaroo Member

"

~. ~

~

!-

Z

Unit F

UJ ~_ _..i

~'~

i ~" i ---~i

~-_

[]

Cordylodus ~pr~176

I

Unit C Unit B

z

>

"

~

Hirsutodontus simplex

Z 0 co

Cordylodus proavus

ch

H. discretus Hispidodontus appressus

~E

I

ii

Unit A

rr"

o

Z <

Z < -ct-

Z O ~>-

< O

--

Hispidodontus resimus

Teridontus nakamurai

<

Z

1::13

<

.....

~

Unit D I

03

300

Unit E --i.,......=

-1-I--

Z

<

Lily Creek Mbr 400--0

Z <

~

137. u_O O O

i

_~

900J

CONODONTS

270~ Polarity Cordyiodus

90~

..1

~

.<

z

~ e~

0

z

no

zonation

Figure 12.10 Black Mountain section (Australia) across Cambro-Ordovician boundary (after Ripperdan and Kirschvink, 1992).

for the Cretaceous quiet zone and PCRS (Kiaman), there is no evidence of substantial polarity bias in the Phanerozoic. At the present state of knowledge of the polarity structure of the paleomagnetic field, the duration of polarity chrons from Middle Carboniferous to the end of the Cretaceous is similar to the distribution of durations seen in the Cenozoic (Fig. 12.11). In the logarithmic histogram of polarity duration since 330 Ma (Fig. 12.12), polarity chron durations generally lie in the 0.1-1 My interval. The PCRS (Kiaman) and Cretaceous quiet zone (KQZ) fall well away from the bulk of the data. Polarity chrons with durations greater than 2 My are very rare. In Figure 12.13, the number of reversals was counted in 4-My bins and

72 68 64

~:

60

.e

56

"~ I--4

52

~9

48

~9

44

Middle Carboniferous to K/T b o u n d a r y

4o

~

36

~, 32

.,..,

~

28

9

~

24

o

20

~ ~ = z

16 12 8

Cenozoic--~

4 ,

0

,

1

,

2

,

3

,

4

9

5

,

6

9

i

7

8

l

,

91011

a

|

|

,

i

.

.

1213141516171819

Duration (Myr) Figure 12.11 A comparison of duration of polarity chrons for the Cenozoic and the Carboniferous-Cretaceous interval.

12011010090804) cO N

70-

O

a) 60C (3)

t~ 5 0 Z

40Q.

30-

9= t~

e" t~

20-

EN .~- o

E .5

0

lO-

TT

T

T

5 0 4 0 30

20

10

6 5

2

1

MY Figure 1 2 . 1 2 Histogram of duration of polarity chrons (log scale), for 330 Ma to present.

232

12 Triassic and Paleozoic Magnetic Stratigraphy

.--4, v

3-

rr 2 > 1 or" '

Cenozoic

A

Cretaceous

150

200 Jurassic

I

250 Triassic

'

'

1 Permian

300

(Ma)

ICarbonifer~

Figure 12.13 Reversalrate from 330 Ma to the present. The number of reversals was counted in 4-My bins and smoothed with a running mean. smoothed with a running mean. The number of reversals per million years ranges from zero to about 4 My-1 in the late Cenozoic. The average reversal rate, excluding the two superchrons, appears to be about 2 My -1. If the record is considered as a whole it would seem that the geomagnetic field has two states, one in which it reverses at a rate of two or more per million years and another in which it does not reverse at all. The change between these two apparent geomagnetic states has been correlated to activity of mantle plumes which rise rapidly from the D" layer and manifest themselves as large igneous provinces such as Ontong Java Plateau and the Deccan Traps (Courtillot and Besse, 1987; Larson and Olson, 1991). McElhinny (1973) analyzed the polarity record from paleomagnetic studies available at that time and concluded that the ratio of normal to reverse polarity indicated that for long periods of time the geomagnetic field seemed to favor one polarity. He recognized polarity bias in favor of normal polarity in Mesozoic rocks and in favor of reverse polarity in the Paleozoic. Intervals of polarity bias were postulated by Algeo (1996) for the Jurassic and Ordovician based on a global analysis of Phanerozoic paleomagnetic data. The analysis given here shows no such bias in the magnetostratigraphic record. The apparent bias is probably caused by unremoved secondary magnetizations acquired during the Kiaman (PCRS) or during the K Q Z and by inadequacies in amount of high quality magneto~ stratigraphic data, particularly for pre-Late Jurassic time.

Secular Variation and Brunhes Chron Excursions

13.1 Introduction "Magnetic stratigraphy" or "magnetostratigraphy" most commonly refers to polarity reversal stratigraphy. However, directional changes associated with secular variation of the geomagnetic field are an important means of correlation in recent sediments. The technique has been largely restricted to lake sediments due to their high sedimentation rates and lack of bioturbation. At typical deep-sea sedimentation rates of about 1 cm/ky, secular variations are averaged out by bioturbation and/or the remanence acquisition process. However, marine sediments from restricted basins (such as the Mediterranean and Black Sea) with sedimentation rates of several tens of cm/ky yielded some of the earliest secular variation records (e.g., Opdyke et al., 1972; Creer, 1974). The first convincing demonstration that secular variation of the geomagnetic field can be recorded in lake sediments is attributed to Mackereth (1971), who used this technique to provide a time frame for sedimentation in Lake Windermere (UK). He correlated remanence declination changes to observatory records, particularly to the early 19th century westerly declination maximum, and showed that the chronology is consistent with the older 14C ages but not with those influenced by detrital carbon input from recent burning of fossil fuels. This work paved the way for the development of the field of secular variation magnetic stratigraphy, which can provide a time frame in lake sediments and is also our main source of information on the nature of Holocene geomagnetic secular variation. 233

234

13 Secular Variation and Brunhes Chron Excursions

Whereas conventional (polarity reversal) magnetic stratigraphy depends on the recognition of directional changes of remanent magnetization of about 180~ secular variations generally do not exceed a few tens of degrees. Hence sampling and measurement techniques in secular variation studies are particularly critical. The development of a pneumatic coring device suitable for lake sampling (Mackereth, 1958) was an important step toward obtaining lake sediment cores that were both long enough and sufficiently undeformed for secular variation studies. In his pioneering study, Mackereth (1971) used an astatic magnetometer to measure subsamples sliced from the cores. The advent of pass-through fluxgate magnetometers (Molyneux et al., 1972) allowed the declination of remanence of the whole core to be rapidly measured without subsampling, and the more recently available 3-axis pass-through cryogenic magnetometers (Goree and Fuller, 1976; Weeks et al., 1993) allow both declination and inclination to be determined rapidly.

13.2 Sediment Records of Secular Variation Although secular variation records can provide a means of precise correlation, only about 20% of all lakes studied have yielded secular variation records of sufficient resolution to define turning points in declination and inclination which can be used for correlation (Thompson and Oldfield, 1986). Some of the clearest and sharpest records are from Western Europe, such as those from Lake Windermere (Mackereth, 1971; Creer et al., 1972), Lough Neagh (Thompson, 1973), Lake Vuokonjarvi (Stober and Thompson, 1977), Lac de Joux (Creer et al., 1980a), and Loch Lomond (Turner and Thompson, 1979). The better dated British records have been merged to give a well-dated "master curve" of directional changes for the last 10,000 years (Fig. 13.1) (Turner and Thompson, 1981) which can be used as a calibration curve. There is also a large body of data from North America which comes mainly from the Great Lakes (Creer et al., 1976; Mothersill, 1979, 1981; Creer and Tucholka, 1982), from Minnesota (Banerjee et al., 1979; Lund and Banerjee, 1985; Sprowl and Banerjee, 1989), from Oregon (Verosub et al., 1986), from California (Lund et al., 1988; Brandsma et al., 1989), and from British Columbia (Turner, 1987). The detailed 36-125 ka record from Mono Lake (California) (see Lund et al., 1988) records the Mono Lake excursion at about 28 ka and a distinctive waveform in the subsequent declination and inclination record which appears to recur every 2500-3500 years (Fig. 13.2). It was suggested that this distinctive waveform evolved out of the Mono Lake excursion and the persistence of the subsequent

13.2 Sediment Records of Secular Variation

235

AVERAGE TRANSFORMED INCLINATION

AVERAGE TRANSFORMED DECLINATION -10.0o 0 ~-,e . . . . . .

10.0~ v::;;;;;:::

1000

,q,

2000.

m

"~,

0

-10.0~

1ooo

/

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~x

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4000

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3000

41300 iF:x m

5000

6000

5000

~ ' ~ J "=

6000

7000

7000 ip

8000

8000,

9000.

10000

5 ~

.J

9000

{'?

10000

Figure 13.1 British master secular variation curve (after Turner and Thompson, 1981). Transformed declination and inclination averaged over ten cores for 0 - 7 ka and three cores for 7-10 ka.

waveform suggests a long-term memory in the core-dynamo process (Lund, 1989). This hypothesis is supported by the work of Negrini et al. (1994), who recorded a repetitive waveform in the secular variation in lake sediments from Oregon which immediately postdate the Pringle Falls excursion. Comparison of the British Columbia (Mara Lake) record with the composite record from the Great Lakes indicates little correlation for the

236

13 Secular Variation and Brunhes Chron Excursions

INCLINATION - I 0 I0 30 50 70 90 i J t t I I I 11 I I

0

ES

I

i

I

I

DECLINATION -40-20 ttlt

0 20 40 It 1 I

f !

-"ZL !

I "'~'~ l '

i

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i B

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< I

I

.

I

~5 c~

I

8

,

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'

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0

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t

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f

I I

_--16

I

--17

I

--to

I~ I

,

I

Figure 13.2 Secular variation record from Mono Lake sediments. The Mono Lake excursion occurs in the 6-7 m interval and is followed by several proposed repetitions of the secular variation waveform (after Lund et al., 1988; Lund, 1989).

13.2 Sediment Records of Secular Variation

2~7

past 1500 years; however, the correlation between 5000 and 2000 yr B.P. (Fig. 13.3) has been interpreted in terms of a lag of 400 years over the 30 degree longitude difference between the two sites, implying a westward drift of the nondipole field of 0.0785 degree/yr (Turner, 1987). Such interpretations are highly dependent on the precision of available (radiocarbon) age control. The correlation of the record from Fish Lake, Oregon (Verosub et al., 1986) with that from British Columbia is problematic (Fig. 13.3), with an apparent 800-yr difference in the age of corresponding features at the two sites (Turner, 1987). This may be explained by radiocarbon ages that are systematically 800 yr too old at the Oregon site (due to the introduction of radiogenically old detrital carbon) or by a time lag between deposition and stabilization of the postdepositional remanence at this site (Turner,

DECL I NA TI ONS 0

-40 . . . .

40 -40

" d .J o,

I''

<

0

o

4o

C .~."

.~ D

40 -40

I

B

C ~.'..

(D

a

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E

,? "~"

F

I NCL I NA TI ONS 40,

0

l

8040 a

,r r

~

.%

80 40

80

bq.I a ' d,~O. c

d dd..""~ c

"ke

f< f ~. ~ , . ~ e

d

h

<

,,~"i

5, Fish Lake

r

Mara Lake

Great

Lake

Figure 13.3 Declination and inclination records from Fish Lake, Oregon (Verosub

et aL,

1986), Mara Lake, British Columbia (Turner, 1987), and Great Lakes (Creer and Tucholka, 1982) (after Turner, 1987). Inflections in the curves are letter coded to indicate proposed correlations.

~

13 Secular Variation and Brunhes Chron Excursions

1987). This dilemma highlights the critical role of radiocarbon and other age determinations in efforts to establish master curves of secular variation and the problems which can arise when radiocarbon dates are not backed by independent age control. The lake sediment records for the last 10 ky from Minnesota (Lund and Banerjee, 1985) and Fish Lake, Oregon (Verosub et al., 1986) can be convincingly correlated with radiocarbon-dated lava flows and archaeomagnetic records (see Brandsma et al., 1989). Elk Lake in Minnesota has yielded a high-fidelity secular variation record with a varve chronology independent of radiocarbon (Sprowl and Banerjee, 1989). Marine records of secular variation are confined to high sedimentation rate environments such as parts of the Mediterranean or Black Sea (Opdyke et aL, 1972); Creer, 1974) and to coastal hemipelagic environments (Brandsma et al., 1989; Levi and Karlin, 1989) and drift deposits (Lund and Keigwin, 1994) where sedimentation rates are of the order of tens of centimeters per thousand years. Lake sediment accumulation rates are often of the order of 1 m/ky, whereas pelagic marine sedimentation rates are characteristically about 1 cm/ky. The relatively enhanced fidelity of lake sediment records of secular variation is partly due to their higher sedimentation rates and partly due to the general absence of bioturbation. At DSDP Site 480 in the Gulf of California, the estimated sedimentation rate is 1 m/ky and the 50-m core was sampled at 10-cm intervals for secular variation studies (Levi and Karlin, 1989). The record cannot be correlated in detail to the lake records although high-amplitude fluctuations in inclination in the 2050 ka interval may be related to low paleointensites at the time of the Mono Lake and Laschamp excursions (Levi and Karlin, 1989). In the Santa Catalina Basin (California Borderlands), estimated sedimentation rates range from 13 to 86 cm/ky (Brandsma et al., 1989). A 3-cm sampling interval in 3-4 m-long piston cores has yielded secular variation records that can be correlated among three cores and with North American lake records (Brandsma et al., 1989), thereby improving the chronology of the sediments from the Santa Catalina Basin for the last 10 ky. Lund and Keigwin (1994) noted that secular variation records from the Bermuda Rise are relatively subdued for the Holocene (where sedimentation rates are about 10 cm/ ky) compared to the Late Pleistocene where sedimentation rates are 2-3 times higher. Secular variation records from Australia (Barton and McElhinny, 1979; Constable and McElhinny, 1985) are more difficult to correlate because of the generally lower amplitude of the directional changes. The records from Victoria (Barton and McElhinny, 1979) cannot be matched with those from Queensland on a swing-by-swing basis (Constable and McElhinny, 1985); however, some features of the VGP paths can be correlated.

13.2 Sediment Records of Secular Variation

~9

The original records from Lake Windermere and Lough Neagh were NRM records, and the observed oscillations in declination and inclination were considered to be periodic with a periodicity of about 2800 years. Partial alternating field demagnetization helped to produce the high-fidelity magnetic record from Loch Lomond (Turner and Thompson, 1979), which, together with an improved time scale, showed that the geomagnetic field changes did not follow a simple oscillationary pattern during the past 7000 years. Turner and Thompson (1981) noted that the calibrated master curve for Britain (Fig. 13.1) has similarities with other European lake and archaeomagnetic records but differs from North American and Japanese records. Creer and Tucholka (1982) considered that the similarities between the UK and North American Great Lakes master curves were sufficient to indicate westward drift throughout most of postglacial time at about the historically observed rate. The general lack of similarity of record from Argentina (Creer et al., 1983b), Australia (Barton and McElhinny, 1979), UK (Turner and Thompson, 1981), and the North American Great Lakes (Creer and Tucholka, 1982) confirms that the secular variation is not due to wobble of the main dipole field axis but to nondipole components with different sources at different sites. For this reason, patterns of secular variation cannot be expected to be correlative over large distances. This is also demonstrated by the lack of similarity of the record from the Sea of Galilee with the UK master curve (Thompson et al., 1985). The Western European record does, however, appear to match over distances of up to about 2000 km, for example, from Western Europe to Iceland (Thompson and Oldfield, 1986). Runcorn (1959) pointed out that westward-drifting nondipole sources will tend to cause the north-seeking geomagnetic vector to precess in a clockwise sense and therefore give a clockwise sense of looping of the VGP paths; however, this interpretation is not unique, and the converse can be true with some source configurations (Dodson, 1979). Presently available European secular variation data indicate a dominantly clockwise sense of looping of the magnetic vectors and of the VGP paths implying that westward drift dominated the past 7000 years or so (with possibly a few hundred years of eastward drift beginning between 1000 and 1500 yr B.P.) preceded by a period of eastward drift (Turner and Thompson, 1981; Thompson and Oldfield, 1986). In the record from British Columbia, clockwise looping from 4.75 ka to 1.5 ka is followed by a period of anticlockwise looping (Turner, 1987). This is broadly consistent with the European record. In the Elk Lake record, clockwise looping predominates with anticlockwise looping in the 500-1500 yr B.P. and 6100-7000 year B.P. intervals (Sprowl and Baner~ee, 1989). The Australian record shows predominantly clockwise looping, with periods of anticlockwise looping between 5.7 and 4.0 ka and

2~0

13 Secular Variation and B r u n h e s Chron Excursions

between 10.5 and 8.8 ka. Within the age constraints, this sense of looping is consistent with the Argentinean record (Constable and McElhinny, 1985). The record from Lac du Bouchet for the 10 ka to 30 ka interval is consistent with predominantly westward-drifting geomagnetic sources (Smith and Creer, 1986). Although there is some consistency between inferred drift direction in the northern hemisphere and in the southern hemisphere, there is no clear global consistency, which has led Creer and Tucholka (1982) to infer that the looping may be largely due to "standing" sources which fluctuate in intensity rather than westward or eastward drift. Crosscorrelation of the Bulgarian archaeomagnetic record (Kovacheva, 1983) with the Elk Lake (Minnesota) record yields a peak at an offset of 520 years, implying a westward drift rate of 0.23~ however, there is little consistency in the drift rate or the drift direction when the same exercise is performed for Eurasian and North American records (Sprowl and Banerjee, 1989). In a recent review of North American Holocene secular variation, Lund (1996) found no evidence for westward or eastward drift. According to Lund (1996), repetitive vector loops may indicate a recurring core generation process manifest in the secular variation record. A global compilation of local changes in declination and inclination for the past 400 years mainly from observatory records (Fig. 13.4) (Thomp-

Virtual

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(1600-1975

AD)

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Historical virtual geomagnetic poles over the E a r t h ' s surface for the 1600-1975 interval. The sense of motion is clockwise except in the region of the Indian O c e a n (after T h o m p s o n and Barraclough, 1982).

13.2 Sediment Records of Secular Variation

~|

son and Barraclough, 1982) illustrates that the amplitude of the secular variation and the timing of distinctive turning points in the records vary from place to place, with the most rapid high-amplitude secular changes and clearest turning points occurring over Africa and Europe. These are therefore the most promising regions for the use of secular variation as a historical correlation tool. Although high-quality relative paleointensity records have been obtained from marine sediments (see Ch. 14), paleointensity studies in lake sediments have generally been less successful. Roberts et al. (1994b) presented a paleointensity record for the 65-105 ka interval from Lake Chewaucan (Oregon) which indicates an NRM/ARM paleointensity high close to 80 ka, consistent with marine paleointensity records (Meynadier et al., 1992; Tric et al., 1992). The paleointensity records from Lake Baikal (Peck et al., 1996) derived from NRM/ARM in three --- 9 m cores yield an excellent match to marine paleointensity records for the last 84 ky. In lake sediments, anhysteretic remanence (ARM) and saturation isothermal remanence (SIRM) are often inadequate as a means of normalizing for variable concentration of remanence-contributing grains, possibly because magnetic grains included in a larger nonmagnetic grains contribute to ARM and SIRM but not to remanence (Turner and Thompson, 1979). Tucker (1980, 1981) has pointed out the difference in nature of the natural DRM and of ARM and has suggested the use of stirred remanent magnetization (StRM) as a normalizing parameter. Thouveny (1987) has demonstrated that StRM can have the same stability against magnetic cleaning as the NRM and that this can be the most suitable normalizing parameter. However, laboratory redeposition of natural samples in known magnetizing fields may not adequately mimic the relationship between NRM and magnetizing field because of the role of organic gels, which appear to be important in stabilizing grain orientation in the natural wet sediment (Stober and Thompson, 1979). Clearly, it would be highly desirable to have an adequate means of determining paleointensities, not only as an additional means of correlation but also to have a complete appreciation of the spatial and temporal nature of geomagnetic secular variation. Within an individual lake, cores can be correlated using tie-points derived from lithological matching and magnetic susceptibility records. A general procedure is to transpose the depth scales for each core to that of a lake master core by linear interpolation between correlation tie-points. In order to correct for variations in tilt of the corer between holes and to find the optimal match between the directional secular variation records from the same lake, several workers (e.g., Constable and McElhinny, 1985; Turner, 1987) have used the cross-correlation technique suggested by Denham (1981). At each step in the iteration, the angle of rotation between two records is calculated to maximize the sum of the scalar products between correspond-

242

13 Secular Variation and Brunhes Chron Excursions

ing unit vectors. In order to use this method, data points of equal depth in the cores must be (linearly) interpolated for each record. The cross-correlation technique applies simple rotations to the whole data set, making no alteration to the depth or time scales. Correction for the tilt of the corer is a critical step as Mackereth-type corers have operated efficiently at tilts up to 35 ~ (Turner, 1987). Matched magnetization vectors from different cores can then be averaged to produce a smoothed master curve for the lake, with estimates of confidence limits using Fisherian statistics. Smoothing of individual records can be achieved by fitting polynomials or splines. Spline functions are generally preferred to polynomials for approximating smooth empirical functions. The commonly used cubic spline function comprises a piecewise polynomial of degree 3 joined at "knots." For a given set of knots, the parameters of the best-fitting cubic spline can be estimated by least squares. The number of knots determines the degree of smoothing and, if equally spaced knots are used, the optimum knot spacing (bandwidth) can be determined by the cross-validation technique (Clark and Thompson, 1978), which involves the repeated comparison of the smoothed record with randomly chosen data points temporally deleted from the record. The optimum bandwidth is that which minimizes the average sum of squares of the differences between excluded observations and the smoothed curve. This process also permits an estimate of the variance of the error measuring the scatter of observations from the smoothed curve, and this is essential for the estimation of confidence limits (Clark and Thompson, 1978). The method can be applied to the smoothing of scalar quantities such as susceptibility or to directional data. Smoothed records can then be correlated by visual matching of distinctive features such as turning points in the directional records. Widely available computer programs such as Analyseries and Corepack are useful for optimizing core correlation by matching distinctive features in the records. A less subjective method of correlation of between-lake records involves determining the stretching function (transformation of depth scales) which results in the optimal fit between records (Clark and Thompson, 1979). The stretching function may be nonlinear, and the method can give valid confidence limits.

13.3 Geomagnetic Excursions in the Brunhes Chron There is a large number of papers in the literature advocating geomagnetic "excursions" (Opdyke, 1972) within the Brunhes Chron, and the evidence for their existence is controversial. Some have suggested that a full reversal of the Earth's magnetic field has occurred many times within the Brunhes and that the excursions should be classified as subchrons (Champion et al.,

243

13.3 Geomagnetic Excursions in the Brunhes Chron

1988). These features may be either extreme excursions in secular variation, aborted reversals of the geomagnetic field, or short polarity subchrons with duration ---104 years. In general, the excursions have proved to be too short in duration and difficult to detect to be particularly useful for stratigraphic correlation purposes, but they are potentially very important for our understanding the nature of the geomagnetic field. It is clear that some of the proposed geomagnetic excursions are based on spurious data. For example, the Gothenberg excursion originated from studies of Swedith varved sediments (Morner et al., 1971). Attempts to replicate the results in sediments from nearby lakes were unsuccessful (Thompson and Berglund, 1976). Both the Gothenberg excursion and the Erieu excursion (Creer et al., 1976) can be attributed to the effects of bottom currents and abrupt facies changes (Thompson and Berglund, 1976; Banerjee et al., 1979). Physical disturbance of sediments, inversion of core segments, and self-reversal mechanisms in igneous rocks may all yield results that could be erroneously interpreted as geomagnetic field behavior. There is a major problem associated with the recognition of geomagnetic events which may have very short duration, local manifestation, and differing character from one place to another. Inability to duplicate the record of such geomagnetic events may not mean that the events do not exist; however, the lack of duplication casts doubt upon the validity of the original record. Nonetheless, four Brunhes Chron excursions have been documented at a number of locations and are now generally considered to represent perturbations of the geomagnetic field (Table 13.1). M o n o L a k e Excursion: The Mono Lake excursion, first documented by Denham and Cox (1971) and subsequently investigated by Liddicoat and Coe (1979), is arguably the best documented excursion of the geomagnetic field during the Brunhes Chron. The excursion (dated at 27-28 ka) is followed by four successive recurrences of the excursion waveform (Fig. Table 13.1 Confirmed Excursions Excursion

Age

Mono Lake

27-28,000 years B.P.

Laschamp

42,000years B.P.

Blake

108-112,000 years B.P.

Pringle Falls Big Lost

218,000 _ 10,000 years B.P. 565,000 years B.P.

References Denham and Cox (1971); Liddicoat and Coe (1979) Bonhomet and Babkine (1967); Levi et al. (1990) Smith and Foster (1969; Tric et al. (1991b); Zhu et al. (1994) Herrero-Bervera et al. (1989, 1994) Negrini et al. (1987); Champion et al. (1988)

~1~

13 Secular Variation and Brunhes Chron Excursions

13.2) (Lund et al., 1988). The directional changes have been closely duplicated at four sites on the shores of Mono Lake, although some sites at this locality with identical tephrostratigraphy failed to record the excursion, exemplifying the problem associated with the recognition of such events. The excursion is recorded as a declination swing to the west by 60 ~ followed by a swing to the east by about 40 ~ and a swing in inclination to - 3 0 ~ (Liddicoat and Coe, 1979). The duration of the excursion has been estimated to be about 10 3 years. Lack of recognition of the Mono Lake excursion at Clear Lake (Verosub, 1977b) or Pyramid Lake (Verosub et al., 1980), located about 300 and 200 km from Mono Lake, respectively, has been used as evidence that the Mono Lake record is spurious (Verosub et al., 1980). These authors considered contemporaneous nondeposition or erosion at Clear Lake and Mono Lake to be highly unlikely. This argument depends critically on the duration of the event; clearly, if the duration was short, recording of the excursion may be due to fortuitously rapid accumulation rates at Mono Lake at this precise time. Two lake records appear to show geomagnetic excursions of comparable age to that recorded at Mono Lake. These records are from Lake Tahoe (Palmer et al., 1979) and from Summer Lake (Negrini et al., 1984), which are 1000 and 600 km from Mono Lake, respectively. Although there are no good age constraints or demagnetization data on the Lake Tahoe record, the Summer Lake excursion is precisely dated by tephrostratigraphy and the magnetic record is well documented. The character of the excursion at Summer Lake is similar to part of that recorded at Mono Lake. The fact that the entire excursion is apparently not recorded at Summer Lake and that the record of the excursion has proved so elusive in lakes with a similar sedimentation history indicates that the excursion has very short duration, probably less than 10 3 years. This in turn leads us to the conclusion that this excursion will be of little use as a means of correlation, although it is well dated both at Mono Lake and Summer Lake by tephrostratigraphy as being in the 24-29 ka range (Liddicoat, 1992; Negrini et al., 1984). The fact that this age correlates closely with the 27-30 ka age for the Lake Mungo excursion (Barbetti and McElhinny, 1976) is an interesting coincidence, but most probably fortuitous. L a s c h a m p Excursion: This excursion was originally detected by Bonhommet and Babkine (1967) in recent lavas of the Puy de Dome region of the Massif Central, France (see also Bonhommet and Zahringer, 1969). The best estimate for its age seems to be 42 ka (Condomes et al., 1982). Detailed studies on these lavas by Heller and Petersen (1982) indicate that they possess self-reversing properties; therefore, the existence of the excursion may be in doubt. The observed directions of magnetization make large angles with the present geomagnetic field direction but are not fully

13.3 GeomagneticExcursions in the Brunhes Chron

24~

reverse polarity. Coring in nearby lake sediments which span the same time interval failed to detect the excursion (Thouveny and Creer, 1992). Recent studies of the Lonchadiere volcanic flow from central France yielded deviating magnetization directions and the lava is the same age as the basalts at Laschamp and Olby. Icelandic lava flows which yield ages indistinguishable from those for the Laschamp excursion give paleomagnetic results which are in accord with the Laschamp excursion (Kristjansson and Gundmundsson, 1980; Levi et al., 1990). Based on these results, the geomagnetic excursion appears to be confirmed. Blake Excursion: Another geomagnetic excursion within the Brunhes Chron which has been well documented is the so-called Blake Event. This excursion was first recorded in the Atlantic (Smith and Foster, 1969) and subsequently both in the Atlantic and in the Caribbean (Denham, 1976). This excursion appears to be a complete, albeit short-lived, reversal of the geomagnetic field and therefore may qualify as a subchron. The appearance of the Blake Event in the sedimentary record is variable, depending, presumably, on the completeness of the record rather than vagaries in the remanence acquisition process. It appears as a single reversal (Tucholka et al., 1987) and two reverse intervals (Creer et al., 1980b; Tric et al., 1991b) in Mediterranean cores and as three more complete reversals in Atlantic cores (Denham, 1976) and in Chinese loess (Fig. 13.5) (Zhu et aL, 1994). The Blake Event has now been recorded in China and in the Atlantic and Mediterranean and therefore appears to be a global event. Age estimates for the Blake Event from Atlantic cores (100 ka, Denham, 1976) compare closely to age estimates from Chinese loess (111-117 ka, Zhu et al., 1994). In Mediterranean piston cores, negative inclinations correlate to oxygen isotopic substages 5e and 5d (Tucholka et al., 1987; Tric et al., 1991b). Tephrochronology and 6180 stratigraphy in these cores indicate fully reverse magnetization directions over a 40-cm interval which the authors estimate to represent ---4000 years. The event is centered at about 110 ka, close to the age originally suggested by Smith and Foster (1969), and best estimates of the duration of the event lie in the 4-6 ky range. Pringle Falls Excursion: The three partially reverse polarity intervals recorded in lake sediments from near Pringle Falls, Oregon (Herrero, Bervera et al., 1989; Herrero-Bervera and Helsley, 1993) were first thought to be a record of the Blake Event. The excursion at Pringle Falls has now been correlated to similar directional records of field excursions at Summer Lake (Oregon) and Long Valley (California), and 4~ dating and tephrochronology indicate an age of 2 1 8 _ 10 ka (Fig. 13.6) (HerreroBervera et al., 1994). Anomalous magnetization directions from the Mamaku ignimbrite in New Zealand (Shane et al., 1994) yield virtual geomagnetic poles which lie close to the VGP path from Oregon/California records

246

13 Secular Variation and Brunhes Chron Excursions

1.Ul,,C~ ~ . . . . /~

NRM (300~

/ K

~

0

I

S u s c e p t i b i l i t y (K) (10 .6 SI) 5 90E

0 ~ I n c l i n a t i o n (o)

I

I

271-9 _

-90

1280 (72 ka)

epth (cm) 1480 (95 ka)

1680 (113 ka)

1880 (130 ka)

Figure 13.5 Record of the Blake Event from Chinese loess (after Zhu et al., 1994).

of the Pringle Falls excursion. Isothermal plateau fission track age on glass (230 ___12 ka) (Shane et al., 1994) and the 4~ age on plagioclase (223 ___3 ka) (McWilliams, personal communication, 1995) from the Mamaku ignimbrite are consistent with the estimated age of the Pringle Falls excursion in Oregon/California. Big L o s t Excursion: Champion et al. (1988) have identified reverse directions in bore holes and outcrops from Idaho dated at 565 _ 14 ka. They called this event the Big Lost subchron. The duration of this possible excursion (or subchron) is not known. A reverse interval recorded in Humbolt Canyon (Negrini et al., 1987; Champion et al., 1988) and low-latitude poles from volcanic rocks in southwest Germany dated at 510 _ 30 ka have been correlated to this excursion (Schnepp and Hradetzky, 1994). Apart from those listed above, Champion et al. (1988) have postulated the existence of five additional subchrons within the Brunhes based on data from the Lake Biwa cores (Kawai, 1984) and from old piston cores (e.g., Goodell and Watkins, 1968; Stearwald et al., 1968; Ryan and Flood, 1972; Watkins, 1968), most of which should not be used for this purpose. Early piston core data containing anomalous magnetization directions (e.g., Ninkovich et al., 1966) were reinterpreted by Wollin et al. (1977) as recording field-related behavior. The cores in question had sedimentation rates of 1 cm/ky or less and it is, therefore, very unlikely that short-term geomagnetic field behavior would have been recorded.

13.3 Geomagnetic Excursions in the Brunhes Chron

DECL I NA TI ON

247'

I NCL I NA TI ON

LongValley

270

180

~.

./

I

h"

~

~'\./

>~

-90

18~

Q

~ '." 0

, '~

I

I

850

.,,, "',-.,,"~: ,..~" i

!

!

0

150

Falls

~a_

j,

::',. :. . . . . . . .

9

.ao '~,

I

\

=

0 ;;""

!

~oii ? 'i ......: .=.:,~,t,.,

; I :I,

-30 -60 t -90

901

9 ,-,., .

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-90

150

Pringle

z,

.

!

I

ri

2,0T

-~o.

'~..:

\

Summer

%

~ ""

"

.,

800

I

Depth (cm)

0

...

,

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- 9 0 / 850

~ / I

,

800

I~

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105~0

~"

~, i

-~o~

";"..... "% i.

-~o/

I "~k

I~

,

,

Depth

o

, 1050

(cm)

Figure 13.6 Declination and inclination of the magnetization in sediments from California and Oregon which record the Pringle Falls geomagnetic excursion, with the correlation of tephra among the sections (after Herrero-Bervera e t al., 1994). Tephras are letter coded to aid correlation from the inclination to the declination records.

In recent years, abundant data have become available from Brunhes age sediments with rates of sedimentation above 1 cm/ky from both conventional piston cores and HPC/APC cores of the Ocean Drilling Program. High-quality data are also available from loess sections of China. In Figure 13.7, these are plotted as straight-line segments assigning the top of the cores to zero age and the Brunhes/Matuyama boundary to 780 ka. In at least one study of loess in China (Heller and Liu, 1984), dense sampling was carried out in an attempt to detect the Blake and Laschamp subchrons; however, no deviating directions were observed. The core with the highest rate of sedimentation is from ODP Leg 107 (Site 650) in the Tyrrhenian Sea, which has a sedimentation rate of 56 cm/ky (Channell and Torii, 1990). Two zones of negative (upward) magnetic inclination are seen in this core;

248

13 Secular Variation and Brunhes Chron Excursions

B

co

~" (~

BL-*

~

'

/

~

, _

/

500 Em.m

I~

4oo

/

8i.~

I=

4oo, 0

1"0

2"0'

30

4"0

50

, aoo. 60

70

., 200 . . . . 80

90

100

S t r a t i g r a p h i c d e p t h (m) Figure

13.7

Positions of reverse polarity intervals in Brunhes age sediment cores and loess

sections (L1 and L2). Brunhes sections range in length from 8 to 400 m. Site 650 from the Tyrrhenian Sea is plotted from right to left with an independent depth scale increased by a factor of 4. The positions of the reverse polarity samples or intervals are indicated by the arrows or in the case of GT (Italy) by a gap in the straight line. It can be seen that a correlation of reverse polarity intervals is not compelling. The positions in time of the Mono Lake (Mo), Laschamp (L), Blake (BL), Jamaica (Ja), Levantine (Lv), Biwa III (BiIII), Emperor (Em), Big Lost (BL), and Delta excursions are shown.

the first (A) occurs at a d e p t h of 10.39 m to 12.31 m and the s e c o n d occurs at a d e p t h of 100.16 to 100.45 m. W h e n p l o t t e d on a time versus d e p t h plot, e v e n t A w o u l d s e e m to c o r r e s p o n d to the M o n o L a k e excursion. E v e n t B occurs at an e x t r a p o l a t e d age of a b o u t 175-ky, an a p p a r e n t age o l d e r t h a n e x p e c t e d for the B l a k e excursion. H o l e 642B (Fig. 13.7) has zones of negative inclination which a p p e a r to correlate with the M o n o L a k e and L a s c h a m p excursions. T h e B l a k e E v e n t , on the o t h e r hand, does not a p p e a r to c o r r e l a t e with negative inclinations in any of the cores and sections e x a m i n e d . This is also true of the Big Lost excursion. Studies of reversal transitions indicate that the length of time t a k e n for a transition to occur is - 5 - 1 0 ky. T h e shortest s u b c h r o n d u r a t i o n you could expect in which the field goes c o m p l e t e l y reverse w o u l d be f r o m 10 to 20 ky. This is, in fact, the length of time e s t i m a t e d for the C o b b M o u n t a i n and R 6 u n i o n subchrons. If full reversals of the m a g n e t i c field can take as

13.3 GeomagneticExcursions in the Brunhes Chron

249

little as 1 ky, as seems to be indicated from the results by Tric et al. (1991b), then very short polarity subchrons may occur. For correlative purposes, however, such short subchrons would seem to be of marginal value except in regions of very high sedimentation rates. Excursion records and records of short polarity events such as the Cobb Mountain subchron (see Clement and Martinson, 1992) imply that the geomagnetic field exhibits a complete spectrum of behavior from high-amplitude secular variation to full polarity reversals. Brief polarity reversals which never stabilize, where the dipole field immediately reverts back to its initial polarity, have also been recorded. Detailed high-fidelity records of the process of polarity reversal, recorded in sediments with high accumulation rates, have occasionally indicated that the dipole configuration of the field is maintained during the reversal process for certain polarity transitions (see Clement, 1992; Clement et al., 1995) and that the paleomagnetic pole follows preferred longitudinal paths during the reversal process (Clement, 1991; Laj et al., 1991) (Ch. 2).

Rock Magnetic Stratigraphy and Paleointensities

14.1 Introduction Rock magnetic stratigraphy involves the utilization of magnetic properties of sediments and sedimentary rocks as a means of (1) stratigraphic correlation, (2) identification of sediment sources and transport mechanisms, (3) characterization/detection of paleoenvironmental change. The first rock magnetic parameter to be widely used for correlation was initial (low field) magnetic susceptibility (k). Its usefulness has long been recognized particularly for correlating volcanoclastic-rich horizons in marine sediments (e.g., Radhakrishnamurty et al., 1968), and it is now a standard correlation technique in deep-sea sediments (e.g., Robinson et al., 1995), lake sediments (e.g., Thompson et al., 1975; Peck et al., 1994), and in loess (e.g., Kukla et al., 1988). Magnetic susceptibility essentially monitors the concentration of ferrimagnetic minerals such as magnetite and maghemite, and therefore the correlation method is sensitive in sediments which show variations in concentration of either of these magnetic minerals. Magnetic susceptibility measurements are part of the standard core scanning procedure on R/V Joides Resolution and other oceanographic research vessels. High-resolution paleoceanographic studies require not only multiple cores at each site (ensuring complete recovery of the sediment sequence) but also the ability to make high-resolution correlation among cores. An important step forward in shipboard core correlation was accomplished during ODP Leg 138 (Lyle et aL, 1992; Hagelberg et aL, 1992). Specialized computer software linked to core scanning systems recording saturated 250

14.2 Magnetic Parameters Sensitive to Concentration, Grain Size, and Mineralogy

25 |

bulk (GRAPE) density, magnetic susceptibility, and digital color reflectance spectroscopy produced optimal correlation of physical parameters between cores. This was the basis for "real-time" core correlation within site, the construction of a composite section for the site, and correlation to downhole logs. Real time within site correlation can be used to determine the number of holes necessary to ensure complete recovery at a site and to compensate for the poorly understood phenomenon of core expansion (often up to 10-20%) during recovery. The field of rock magnetic stratigraphy has broadened considerably in the last 15 years with the use of magnetic parameters which are sensitive not only to magnetic mineral concentration but also to the grain size and mineralogy. These techniques were largely developed to study environmental changes in lake catchments by monitoring changes in lake sediment magnetic mineralogy which are sensitive to changes in land use, erosion, industrialization, and climate (see Thompson and Oldfield, 1986). Similar methods have been applied to marine sediments (Robinson, 1986; Bloemendal et aL, 1988; Doh et al., 1988) and loess deposits (Maher and Thompson, 1991, 1992) and have been shown to provide not only a sensitive means of correlation but also information on changing detrital and biogenic fluxes. The magnetic parameters useful for correlation and/or environmental modeling depend on the nature of the variations in magnetic mineralogy within the sediment.

14.2 Magnetic Parameters Sensitive to Concentration, Grain

Size, and Mineralogy

The magnetic parameters listed below and in Table 14.1 are the principal parameters used in studies of marine and lake sediments (see Thompson et al., 1975; King et al., 1982; Thompson and Oldfield, 1986; Robinson, 1986; Bloemendal et al., 1988; Doh et al., 1988; King and Channell, 1991; Stoner et al., 1996).

a. Magnetic Susceptibility (k) Volume (initial) magnetic susceptibility, defined as volume magnetization divided by applied field, is often considered to be a measure of ferrimagnetic mineral concentration; however, susceptibility (in magnetite) shows a grain size dependence (see Maher, 1988) increasing through the SDPSD-MD range. As the susceptibility of the ferrimagnetic minerals magnetite and maghemite is 3 or 4 orders of magnitude greater than that of common antiferromagnetic minerals, such as hematite and goethite,

252

14 Rock Magnetic Stratigraphy and Paleointensities

the susceptibility is usually dominated by the ferrimagnetic grains and therefore gives information on the presence or absence of this group of magnetic minerals. For magnetite- and maghemite-bearing sediments, magnetic susceptibility is a reasonable measure of the concentration of the magnetic mineral, and is not in itself a sensitive measure of grain size.

b. Anhysteretic Remanent Magnetization (ARM) or Anhysteretic Susceptibility (ka.n) and Saturation Remanent Magnetization (SIRM) These parameters are also concentration dependent and, as for susceptibility, are dominated by the ferrimagnetic grains present in the sample. These parameters are particularly useful because their magnitude tends to increase with decreasing grain size of the ferrimagnetic grains (see Maher, 1988). Normalizing these parameters by initial susceptibility partially compensates for variations in concentration and hence the SIRM/k ratio (Thompson, 1986) and ARM/k o r karm/k ratio (King et aL, 1982) vary inversely with magnetic particle size and are therefore useful granulometric parameters. The former is likely to be more sensitive to the multidomain grain size range and the latter to the pseudo-single-domain range. The granulometric method utilizes the fact that ARM and SIRM decrease with increasing grain size whereas susceptibility is relatively insensitive, showing slight increase with increasing grain size. Interpretation is complicated by the nonlinear response of these parameters to changes in grain size. An additional drawback of all three granulometric ratios is that ultrafine superparamagnetic (SP) grains will have negligible remanence (SIRM, ARM, karm) but relatively high k, resulting in low values of the ratios (SIRM/k ARM/k, or karm/k), comparable to those seen in coarse multidomain grains.

c. S-Ratio and "Hard" IRM (HIRM) The S-ratio (Stober and Thompson, 1979; Bloemendal, 1983) is found by measuring the SIRM and then placing the sample in a back field (typically 0.1 T) and remeasuring the IRM. S_o. 1 =

(IRM-o.1/SIRM)

S-0.1 is a convenient measure of the proportion of (coarse) low-coercivity ferrimagnetic grains to higher coercivity grains. As the proportion of low-coercivity grains saturating in fields less than 0.1 T increases, this Sratio approaches unity. For discriminating ferrimagnetic grains (such as magnetite) from high-coercivity antiferromagnetic grains such as geothite

14.3 Rock Magnetic Stratigraphy in Marine Sediments

253

or hematite, S-0.3 (determined using a back field of 0.3 T) would be more appropriate. The HIRM ("hard" IRM) is the difference between the IRM-0.3 (IRM in a 0.3-T field) and SIRM divided by 2 (Robinson, 1986; Bloemendal et al., 1988). HIRM-0.3 = (SIRM + IRM_0.3)/2 This parameter is also a measure of the concentration of antiferromagnetic minerals with high coercivity. The two parameters (S-ratio and HIRM) complement each other and might be expected to show an inverse relationship in cores exhibiting a wide range of magnetic mineralogy and grain size. The combination of these two parameters provides a means of monitoring changes in magnetic mineralogy and magnetic mineral concentrations. The coercivity range in which the S-ratio and HIRM are sensitive will depend not only on the value of the back field but also the magnetizing field used to acquire the SIRM. Note that, in practice, the "SIRM" is often acquired in a magnetizing field of about 1 T, which is close to the maximum field attainable by most small electromagnets and pulse magnetizers. This field may not fully saturate the sample, particularly in the presence of high-coercivity antiferromagnetic minerals such as goethite or hematite.

d. Frequency-Dependent Susceptibility The susceptibility of fine-grained magnetite, with grain sizes close to the superparamagnetic/single domain boundary, is frequency dependent, decreasing with increasing frequency of the applied field (Stephenson, 1971). The susceptibility of single-domain and multidomain grains is relatively unaffected by changes in frequency. The frequency-dependent susceptibility (kfd) is often expressed as a percentage, where kfd = (khf--k~f)/k~f and khf and k~f are the susceptibility in the high-frequency (typically about 5-10 kHz) and low-frequency field (typically about 1 kHz), respectively. The frequency-dependent susceptibility is a convenient measure of the concentration of very fine grained (SP) magnetite in the sample.

14.3 Rock Magnetic Stratigraphy in Marine Sediments Over 15 years ago, the recognition of orbital frequencies in NRM sediment records led to the proposal that geomagnetic field intensity was orbitally modulated (e.g., Wollin et al., 1977). It has since been demonstrated that these observations are best explained by orbitally controlled climatic varia-

254

14 Rock Magnetic Stratigraphy and Paleointensities

Table 14.1 Generalized Table of Downcore Magnetic Parameters and Their Interpretations Parameter

Interpretation

Bulk Magnetic Measurements Dependent on mineralogy, concentration, and Natural Remanent Magnetization grain size of the magnetic material as well as (NRM): The fossil (remanent) mode of acquisition of remanence and magnetization preserved within the sediment. NRM recorded as intensity and direction of the geomagnetic declination, inclination, and intensity. field. Volumetric Magnetic Susceptibility (k): k is a first-order measure of the amount of A measure of the concentration of ferrimagnetic material (e.g., magnetite): k is particularly enhanced by superparamagnetic magnetizable material. Defined as the ratio of induced magnetization (SP) magnetite (<0.03/~m) and by large intensity (M) per volume to the magnetite grains (>10 p,m). When the strength of the applied weak field concentration of ferrimagnetic material is (H): k = M/H. low, k responds to antiferromagnetic (e.g., hematite), paramagnetic (e.g., Fe, Mg silicates), and diamagnetic material (e.g., calcium carbonate, silica) which may complicate the interpretation. Isothermal Remanent Magnetization SIRM primarily depends on the concentration (IRM): Magnetic remanence acquired of magnetic, principally ferrimagnetic, under the influence of a strong DC material. It is grain size dependent, being particularly sensitive to magnetite grains field. Commonly expressed as a smaller than a few tens of microns. saturation IRM or SIRM when a field greater than 1 T is used. A back field IRM (BIRM) is that acquired in a reversed DC field after SIRM acquisition. k ARM is primarily a measure of the Anhysteretic Remanent Magnetization concentration of ferrimagnetic material: (ARM): Magnetization acquired in a however, it is also strongly grain size biasing DC field within a decreasing d e p e n d e n t , k ARM preferentially responds to alternating field. Commonly smaller magnetite grain sizes (<10/~m) and expressed as anhysteretic is useful in the development of grain size susceptibility (kARM)when normalized by the biasing field dependent ratios. used. Constructed Magnetic Parameters HIRM is a measure of the concentration of The "hard" IRM (HIRM): This is magnetic material with higher coercivity than derived by imparting a back field, the back field. This commonly gives typically 0.1 or 3.0 T, on a sample information on the concentration of previously given a SIRM. The antiferromagnetic (e.g., hematite) or very resulting B IRM, which has a negative fine grained ferrimagnetic (e.g., magnetite) sign, is used to derive the HIRM by grains depending on the back field used. the formula: HIRM = (SIRM + BIRM)/2.

14.3 Rock Magnetic Stratigraphy in Marine Sediments

255

Table 14.1 continued Parameter

Interpretation

S ratios: These are derived by imparting a back field, typically 0.1 or 0.3 T, on a sample previously given a SIRM. The resulting B IRM, which has a negative sign, is normalized by the SIRM; S = BIRM/SIRM. This provides a measure of the proportion of saturation at the back field applied.

The S ratios can be used to estimate the magnetic mineralogy (e.g., magnetite or hematite). Downcore variations may be associated with changing mineralogy. Values close to -1 indicate lower coercivity and a ferrimagnetic mineralogy (e.g., magnetite): values closer to zero indicate a higher coercivity, possibly an antiferromagnetic (e.g., hematite) mineralogy. S ratio with a back field of 0.1 T may be sensitive to mineralogic and grain-size changes, whereas the S ratio with the 0.3 T back field is more sensitive to mineralogical changes (e.g., proportion of magnetite to hematite). If the magnetic mineralogy is dominantly magnetite, kARM/k varies inversely with magnetic grain size, particularly in the 110 tzm grain-size range. However, the interpretation of this ratio may be complicated by significant amounts of superparamagnetic (SP) or paramagnetic material. If the magnetic mineralogy is dominantly magnetite, SIRM/k varies inversely with magnetic particle size. SIRM/k is more sensitive than k ARM/k to changes in the proportion of large (>10/xm) grains. SIRM/ k may also be compromised by SP or paramagnetic material. SIRM/kARM increases with increasing magnetic grain size but is less sensitive and can be more difficult to interpret than the two ratios above. A major advantage of SIRM/kARM is that it only responds to remanence carrying magnetic material and is therefore not affected by SP or paramagnetic material. kf is used to indicate the presence of SP material. SP material in high concentrations can compromise the grain-size interpretation made using k ARM/kand SIRM/k.

k ARM/k: Indicates changes in magnetic grain size, if the magnetic mineralogy is dominantly magnetite.

SIRM/k: Indicates changes in magnetic grain size, if the magnetic mineralogy is dominantly magnetite.

SIRM/kARM: Indicates changes in magnetic grain size, if the magnetic mineralogy is dominantly magnetite.

Frequency Dependent Magnetic Susceptibility (ky): The ratio of lowfrequency k (0.47 kHz) to high frequency khf (4.7 kHz) calculated by kf = l O 0 • khf)/k.

Hysteresis Measurements a Mrs~Msdecreases with increasing magnetite

Saturation Magnetization (Ms): Ms is the magnetization within a saturating field; Saturation Remanence (Mrs): Mrs is the remanence remaining after removal of the saturating field.

grain size in the submicron to few tens of microns grain size range.

(continues)

2~6

14 Rock Magnetic Stratigraphy and Paleointensities

Table 14.1 continued

Parameter Coercive Force (He): The back field required to rotate saturation magnetization to zero within an applied field; Coercivity of Remanence (Hcr): The back field required to rotate saturation magnetization to zero remanence.

Interpretation For magnetite, the ratio ncr/Hc increases with increasing grain size (in the submicron to several hundred micron grain-size range) due to the strong grain-size dependence of both parameters, particularly Hc. ncr is a useful guide to magnetic mineralogy.

Note: After Stoner et al., 1996. a Hysteresis parameters provide a means of monitoring grain-size variations in magnetite. Sediments should be homogeneous because of the small sample size (<0.05 g) typically used for hysteresis measurements. Mixed magnetic mineral assemblages greatly complicate the interpretations. The generalized interpretations listed above are based on a ferrimagnetic (e.g., magnetite) mineral assemblage.

tions affecting magnetic mineral concentrations, which in turn affect N R M intensities (Amerigian, 1974; Chave and Denham, 1979; Kent, 1982). In deep-sea sediments, the principal climatically controlled variable which influences NRM intensity is carbonate content, which dilutes the magnetic mineral flux (Kent, 1982). Detailed studies of the magnetic mineralogy of marine sediment cores (Robinson, 1986; Bloemendal et al., 1988; Doh et al., 1988) have shown that the climatically induced carbonate dilution of magnetic mineral concentrations is not the only paleoclimatic process influencing the magnetic properties and that changes in eolian, biogenic, and ice-rafted fluxes of magnetic minerals can be deduced from changes in magnetic properties. Susceptibility records at ODP sites off west Africa and in the Arabian Sea are related to terrigeneous (eolian) flux (Bloemendal and deMenocal, 1989). Spectral analysis of these Late Pliocene/Pleistocene susceptibility records reveal 100 ky, 41 ky, 23 ky, and 19 ky periodicities (Fig. 14.1). There is an apparent shift in power at about 2.4 Ma, coincident with the onset of Northern Hemisphere Glaciation. Prior to this time, the 23 ky and 19 ky (precession) periodicities dominate, and after this time the 41 ky (obliquity) periodicity dominates. The presence of the 41 ky periodicity during the Matuyama Chron is also a feature of the North Atlantic ~180 record (Raymo et al., 1989). The shift in power in the off-Africa susceptibility records implies a change in monsoon intensity and/or eolian source area climate related to the onset of Northern Hemisphere Glaciation (Bloemendal and deMenocal, 1989). A similar change in power at this time (---2.4 Ma) has been noted in foraminiferal abundance data from the Mediterranean (Sprovieri, 1992).

257

14.3 Rock Magnetic Stratigraphy in Marine Sediments

721 Susc. 661 Susc. cycles kyr -1 (IJ SI Units) cycles kyr -1 (~J SI Units) 0 350 0 0.08 nO . . . . . . . . 0 0.08

~i

o. i 1.0

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41 19 Periodicities (Ky) .=riodicities (Ky)

Figure 14.1 Susceptibility records from the central Atlantic (Site 661) and Arabian Sea (Site 721) coasts of Africa. Spectral analysis over 0.4 My intervals indicates power in the Milankovitch periodicities (after Bloemendal and deMenocal, 1989).

In the King's Trough Flank area of the North Atlantic (41-43~ peaks in susceptibility (k) tend to coincide with glacial intervals (Fig. 14.2) as indicated by the oxygen isotope record, cold water foraminiferal assemblages, and lows in calcium carbonate concentrations (Robinson, 1986; Robinson et al., 1995). In core BOFS-5K, located further north at 50.6~ (Fig. 14.2), the susceptibility record is dominated by the ice-rafted detritus (IRD) associated with Heinrich events, as observed in other piston cores from this region (Grousset et al., 1993). Robinson et al. (1995) have used the ratio of susceptibility at the Last Glacial Maximum (LGM) (---1819 ka) to susceptibility in the Holocene as a means of mapping the distribu-

Figure 14.2 Susceptibility records from North Atlantic piston cores illustrating the change in the record from the carbonate content-controlled record in the King’s Trough area (-41-43”N) to the region close to 50”N where Heinrich (IRD) events dominate the record (after Robinson et al., 1995).

14.3 Rock Magnetic Stratigraphy in Marine Sediments

259

tion of I R D during the L G M (Fig. 14.3). The ice sheet instability which leads to high-susceptibility Heinrich ( I R D ) events in the North Atlantic (Broecker et al., 1992) has a different manifestation in the Labrador Sea. In this region, detrital carbonate (DC) and low detrital carbonate (LDC) layers, some of which correlate to Heinrich events and other I R D layers in the North Atlantic, have been associated with turbiditic activity in the North Atlantic Mid-Ocean Channel ( N A M O C ) (Stoner et aL, 1996). Some 50 ~

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1 4 . 3 Ratios of magnetic susceptibility at the Last Glacial Maximum (LGM, --+1819 ka) to susceptibility in the Holocene, used to map terrigeneous ice-rafted detritus during the LGM (after Robinson et aL, 1995).

260

14 Rock Magnetic Stratigraphy and Paleointensities

DC and L D C events are characterized by high susceptibility (due to their association with IRD) although the detrital events are more easily recognized by ferrimagnetic grain size-sensitive ratios such as karm/k and SIRM/ k (Fig. 14.4). Such records provide a link between (Laurentide) ice sheet variability and the marine record which are critical to the understanding of climate change during the last glacial cycle. Parameters which are not strictly concentration dependent and therefore independent of calcium carbonate content, such as the S ratio and the

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Figure 14.4 In Labrador Sea piston core P-094, the grain size-sensitive parameters (karm/k and SIRM/k) are more sensitive than susceptibility (k) for recognition of detrital carbonate (DC) and low detrital carbonate (LDC) layers (after Stoner et al., 1995a).

t I I ~

14.3 Rock Magnetic Stratigraphy in Marine Sediments

261

ARM/k and ARM/SIRM ratios, show variations which reflect not only the principal glacial/interglacial variations but also additional fine-scale variations, particularly within interglacials (Fig. 14.5). Robinson (1986) interpreted these data in terms of three magnetic mineral fluxes: (1) antiferromagnetic mineral flux mainly of eolian origin associated with vigorous atmospheric circulation and aridity in glacial periods; (2) coarse ferrimagnetic mineral flux associated with ice rafting in glacial periods; and (3) a constant flux of fine-grained (single-domain) magnetite of eolian or biogenic origin, the effect of which is noticeable when fluxes (1) and (2) are low in interglacial periods. Changes in k and karm/k at Termination I and II for a core located off the southern tip of Greenland are interpreted in terms of progressive deglaciation as the continental ice sheet melted back to the shoreline and then retreated into the continental interior (Stoner et al., 1995a). At Termination I in this core, karm/kperturbations record the Younger Dryas (YD), Heinrich Layer 1 (H1), and an additional melt event (ME1) prior to a broad

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262

14 Rock Magnetic Stratigraphy and Paleointensities

low in the karm/k record (Fig. 14.6). ME1 is interpreted to signify detrital influx at the time of initial detachment of grounded sea ice from the shelf. The subsequent broad low in karm/k is interpreted as the detrital influx during continental deglaciation on Greenland (Stoner et al., 1995a). This

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14.3 Rock Magnetic Stratigraphy in Marine Sediments

2~

interpretation is supported by radiocarbon age control tying the low in karm/k to geomorphological evidence for the timing of continental ice retreat on Greenland (Funder, 1989). Bloemendal et al. (1988) used seven rock magnetic parameters as a means of detailed correlation in Quaternary cores from the eastern central Atlantic. Rock magnetic parameters correlate to oxygen isotope records, with HIRM and k~rm/SIRM showing the strongest correlation with the oxygen isotope curve. Magnetic parameters are computed both as magnetic accumulation rates (computed by dividing the parameter by the estimated time taken for each individual sample to accumulate; Bloemendal, 1983) and as accumulation rates of the noncarbonate content. This allows the variations in magnetic properties independent of total sedimentation rate and carbonate sedimentation rate to be assessed. An antiferromagnetic and coarse ferrimagnetic detrital flux is dominant in glacial intervals, and a fine-grained ferrimagnetic flux is largely confined to warm interglacial events. The principal indicator for this latter component (karm) shows a strong 21-23 My precessional variance in one of the cores. Bivariate scatter plots of karm against k and S-0.3 against karm/k have been used as a means of discriminating the magnetic mineralogy of late Neogene and Pleistocene sediments from different sedimentological environments (Bloemendal et al., 1992). The karm against k plot is largely controlled by the grain size spectrum of ferrimagnetic grains (magnetite and maghemite), and the S-0.3 against karm/k plot by the ratio of antiferromagnetic (hematite) to ferrimagnetic grains. Differing sedimentary environments tend to reveal characteristic distributions in these bivariate plots. For example, eolian detritus is rich in antiferromagnetic grains, ice-rafted detritus is poorly sorted and rich in coarse ferrimagnetic material, and current-controlled deposition tends to be relatively well sorted. Doh et al. (1988) studied a 25-m core (LL44-GPC3) from the central North Pacific which covers the entire Cenozoic. The lithology is the unfossiliferous pelagic clay which is characteristic of the Pacific midlatitudes (Davies and Gorsline, 1976). The stratigraphic control in this facies is particularly poor due not only to the almost total lack of microfossils but also to the unstable nature of the remanent magnetization. The magnetic parameters show coherent fluctuations downcore which are potentially useful for correlation. In addition, fluctuations in concentration-dependent parameters (such as HIRM) plotted as accumulation rates correlate well with the sedimentologically determined total eolian accumulation rate (Fig. 14.7), indicating that these parameters can be used as a proxy indicator for eolian

264

14 Rock Magnetic Stratigraphy and Paleointensities

Eolian Detritus [mg cm'=(10a yr)"~] 4

20 v

j

40 "~

"l

60 -

g

HIRM

(xl03A/m.y.) 0

80 200 400 9

,

,

,

10

Miocene 20

20

3o A m

m

"g

4o

O

40 50

60

90

9

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60 70

>tr. ,a:: Oligocene m I-. rr I..I.I I-I Eocene

Paleocene CRETACEOUS

Figure 14.7 HIRM as a proxy for eolian detrital flux in red clays from the equatorial Pacific (after Doh et aL, 1988).

flux (Doh et aL, 1988). Changes in eolian flux are related to changes in paleogeography and paleoclimate, as the Pacific plate moved northward, and can provide the basis for local stratigraphic correlation; however, the precision of the correlation technique depends critically on whether the magnetic properties are due to changes in detritus or to later diagnetic growth/alteration of magnetic minerals. In core LL44-GPC3, polarity intervals (stable magnetizations) are recorded above 4.33 m (late Pliocene/ Pleistocene) and below 19 m (Paleocene/early Eocene) (Prince et aL, 1980). The intervening interval is characterized by unstable, highly viscous remanence. The intervals of stable remanent magnetization coincide with low values of ARM/k, indicating that increases in grain size of magnetic minerals may be the cause of remanence instability (Doh et al., 1988). In these intervals, magnetic grain size (from ARM/k) and eolian grain size (from grain observation) are high and vary consistently, suggesting that most of the magnetic minerals in these intervals are of eolian origin. This is not the case for the interval of unstable remanence, where eolian grain size fluctuations are not recorded in the ARM/k profile, suggesting diagenetic alteration of magnetic minerals in this interval. Indeed, the instability of the remanent magnetization in this particular facies has been attributed to oxidation of primary magnetite to maghemite downcore (Kent and Lowrie,

14.4 Rock Magnetic Stratigraphy in Loess Deposits

265

1974; H. P. Johnson et al., 1975) or alternatively the growth of authigenic ferromanganese phases during diagenesis (Henshaw and Merrill, 1980). The critical observation is whether the boundary between the stable and the unstable remanence record, which is controlled by the magnetic mineralogy, occurs at a consistent time horizon or stratigraphic level. Available data indicate that it tends to occur very close to the Gauss/Matuyama boundary (Opdyke and Foster, 1970; Prince et al., 1980; Yamazaki and Katsura, 1990) although synchroneity has not been well established. The change in magnetic mineralogy in the late Pliocene may be either directly or indirectly due to change in eolian flux, which in turn may be related to the onset of Northern Hemisphere Glaciation. If so, these mineralogical changes represent time lines suitable for regional stratigraphic correlation.

14.4 Rock Magnetic Stratigraphy in Loess Deposits Heller and Liu (1982) pointed out that the susceptibility variations in the Chinese loess match the paleosol/loess lithologic variability, and Heller and Liu (1984) matched the record to the Brunhes oxygen isotope record from the Pacific of Shackleton and Opdyke (1976). Kukla et al. (1988) showed that the susceptibility record closely parallels the orbitally tuned SPECMAP oxygen isotope record derived from deep-sea sediments (Fig. 14.8), indicating a close linkage between eolian flux, global ice volume, and climate. Rock magnetic data have played an important role in understanding the origin of the climatic record in Chinese loess. The present consensus is that the higher susceptibility of the paleosols in the Chinese loess is the result of "magnetic enhancement" associated with authigenic production of SP and SD magnetite during pedogenesis (Zhou et al., 1990; Heller et al., 1991; Maher and Thompson, 1991, 1992; Banerjee et al., 1993; Evans and Heller, 1994). Beer et aL (1993) showed that the pedogenic and detrital susceptibility components in the loess can be quantitatively distinguished by comparing the susceptibility variations with 1~ concentrations. According to these calculations, the susceptibility in the paleosols due to magnetite formed in situ is often ---50% greater than the detrital susceptibility. Taking presentday precipitation/susceptibility data (Liu et al., 1992) as the means of calibration, Heller et al. (1993) used estimates of the pedogenic susceptibility contribution to estimate paleoprecipitation. Maher et al. (1994) constructed a paleoprecipitation record based on bulk susceptibility differences between paleosols and intervening unweathered loess and a logarithmic relationship between rainfall and susceptibility in nine modern paleosols across the Chinese loess plateau.

266

14 Rock Magnetic Stratigraphy and Paleointensities

ODP 677

Xifeng

~lS0 (o/~ to PDB) -1 ,

Susceptibility (10.8 m~kg) Lithology

-2

9

--

0

,

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'~

200 "

~

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.

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~

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,

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Figure 14.8 Correlation of susceptibility records from Chinese loess with the benthic oxygen isotope record from O D P Site 677 (after Heller and Evans, 1995). Shading in lithology indicates paleosols separated by loess.

The pedogenic model which explains the magnetic variability in Chinese loess cannot be applied to all loess deposits. The susceptibility record in Alaskan loess has orbital variability, but the intervals of high susceptibility correlate to cold climate intervals, opposite to the relationship in the Chinese loess (Beg6t and Hawkins, 1989). A close relationship between mean loess grain size and susceptibility values implied that the high-susceptibility intervals can be attributed to greater wind velocities during glacial intervals, which transported larger magnetite grains which dominate the susceptibility record (Beg6t e t a L , 1990).

14.5 Rock Magnetic Stratigraphy in Lake Sediments Magnetic parameters provide a basis for correlation and for determining changes of magnetic mineral fluxes in both marine and lake sediments. However, because lake sedimentation rates are typically 2 or 3 orders of

14.5 Rock Magnetic Stratigraphy in Lake Sediments

~7

magnitude greater than those of deep-sea sediments, the time scale of the observations is very different. Whereas magnetic mineral flux changes in the marine setting can give us information on long-term climate change, the same information from lake sediments is generally giving us information on Holocene and recent anthropogenic environmental change. There are, however, lake sediment cores which extend back through the last glacial cycle such as those from Lac du Bouchet (Thouveny et al., 1994) and Lake Baikal (Peck et al., 1994). In these studies the susceptibility record is an important means of monitoring lithologic changes associated with the glacial-interglacial cycling. Volume susceptibility measurements are a very important, rapid, and nondestructive means of core correlation in lake sediments (Thompson et al., 1975; Bloemendal et al., 1979; Peck et al., 1994). Distinctive features in the records, which are controlled by variations in concentrations of ferrimagnetic minerals, can be matched to monitor variations in sediment accumulation. The method can be useful even in lakes which show wide variations in sedimentation rate (Thompson and Morton, 1979). Other magnetic parameters, such as SIRM, can also be used for core correlation and may be useful when susceptibility values are uniformly low, but they may require subsampling of the whole-core and therefore measurements may be considerably slower than whole-core susceptibility measurements. The use of magnetic susceptibility profiles for correlation in lake sediments led to the observation, at Loch Neagh, that progressively increasing susceptibility values parallel pollen variations signifying progressive deforestation (Fig. 14.9) (O'Sullivan et al., 1972). Thompson et al. (1975) interpreted these data to indicate increased erosion and a resulting flux of ferrimagnetic grains from the basaltic bedrock. It has since been demonstrated that the susceptibility of topsoil can be greater than that of the deeper soil horizons and of the bedrock. This "magnetic enhancement" has been attributed to the formation of magnetite and maghemite in the surface soil (Ozdemir and Banerjee, 1982; Mullins, 1977; Maher and Taylor, 1988; Evans and Heller, 1994; Eyre and Shaw, 1994). Increases in susceptibility associated with pollen evidence for deforestation would now be interpreted to reflect increased erosion of topsoil as a result of changes in land use (e.g., Dearing et al., 1986; Thompson and Oldfield, 1986). It has been demonstrated that susceptibility of lake sediments can be a very sensitive indicator not only of changes in erosion style in the catchment but also of changes in climate (Thompson and Oldfield, 1986), changes in stream discharge (Dearing and Flower, 1982), natural burning or burning for land clearance (Rummery et al., 1979; Rummery, 1983), and industrial burning of fossil fuels (Oldfield et al., 1983). Burning contributes to the lake sediment susceptibility record because magnetite is an important product of both soil burning (Longworth et aL, 1979) and industrial burning (Chaddha and

268

14 Rock Magnetic Stratigraphy and Paleointensities

k (10 "6 SI units)

o

,

2s

E

15

~2 1.0

Cereal pollen

2.0

I .-....

Pollen Zones

~ ~

~.~ ~

"~ C~C~

-.. /

9 omo

ra~

.~ ~ ~

~

0

~0

,~0

% Grass pollen

Figure 14.9 6000-year record from Lough Neagh (Northern Ireland). Forest clearance and farming are indicated by increases in plantains, grasses, and cereal pollen. Increases in susceptibility are attributed to increased soil erosion in the lake catchment (after Thompson and Oldfield, 1986).

Seehra, 1983). As many of these anthropogenic processes produce changes in ferrimagnetic influx, concentration-dependent parameters such as susceptibility and SIRM are most applicable (rather than mineralogy-dependent magnetic parameters). However, for detailed studies of erosion history and land use changes, a more complete characterization of sediment source material and of the sediment itself is necessary. Magnetic parameters can be used as tracers to determine changes in sediment source and thereby provide markers of environmental and land use change. In this type of

14.6 Future Prospects for Rock Magnetic Stratigraphy

'269

application, there is a need to establish a more formal, quantitative approach to the study of magnetic mixtures, in order to delineate the proportion of individual magnetic constituents. Thompson (1986) has used the Simplex search method to determine optimal mixing models on the basis of particular magnetic characteristics (such as IRM acquisition curves) of individual sediment samples and of possible source materials. Alternatively, the magnetic characteristics of the sediment can be matched to combinations of standard curves based on synthetic samples.

14.6 Future Prospectsfor Rock Magnetic Stratigraphy Until recently, rock magnetic stratigraphy was confined to the use of susceptibility as a convenient method of core correlation in volcanogenic marine sediments. During the past ten years, rock magnetic studies of lake sediments have demonstrated that rock magnetic parameters are not only useful for core correlation but can provide detailed information on changes in magnetic mineral flux that can be related to change in land use (Thompson and Oldfield, 1986). The techniques used in this developing field of environmental magnetism are now being applied to the marine and terrestrial sedimentary record. Susceptibility data from Chinese loess deposits (Kukla et al., 1988) and from the eolian-dominated equatorial sites off Africa (Bloemendal and deMenocal, 1989) have variance in which the dominant "heartbeat" is very similar to that seen in marine oxygen isotope records, providing a fascinating link between marine and terrestrial climate records. Rock magnetic methods are generally conducted on subsamples from sediment cores often collected as 8-cm 3 plastic cubes. The widely used 2G755R cryogenic magnetometer (with 4.2-cm-diameter access) is designed for discrete sample measurement with homogeneous response of SQUID sensing coils over a volume large compared to the discrete sample volume. Weeks et al. (1993) describe a modified 2G-755R magnetometer in which the SQUID sensing coils are arranged for optimal spatial resolution of magnetization from continuous "u-channels" tracked through the sensing region. Deconvolution of the signal in the time domain (Constable and Parker, 1991; Oda and Shibuya, 1996) gives resolution comparable to that achieved by back-to-back discrete sampling. The u-channels (see Tauxe et al., 1983b; Nagy and Valet, 1993) are typically 1.5 m in length with a 2 x 2 cm cross-section. Stepwise demagnetization of NRM, acquisition and stepwise demagnetization of IRM and ARM can be carried out on a single automated measurement track (Weeks et al., 1993). Rock magnetic measurement tracks with this capability will greatly increase the speed at which

270

14 Rock Magnetic Stratigraphy and Paleointensities

such data can be acquired and will result in increased use of these data for stratigraphic and paleoenvironmental studies.

14.7 Paleointensity Determinations Over the 40-year history of modern paleomagnetic research, the vast majority of the studies have dealt with directional variability of the ancient geomagnetic field with very little attention given to paleointensity studies. In recent years, there has been a renewed interest in paleointensity studies for several reasons: (1) Paleointensity information is an important component of understanding the field generation process, the reversal process, and factors controlling reversal frequency. (2) Paleointensity data from sediments appear to record the time-average dipole field intensity, the variations of which may be a useful means of high-resolution global correlation. (3) The atmospheric production of radionuclides such as ~4C and ~~ is partly controlled by geomagnetic field intensity, due to the role of the field in shielding the low-latitude atmosphere from incoming cosmic rays. Paleointensity records provide a potential means of monitoring variations in 14C and 1~ production, improving the potential precision of radiocarbon ages and possibly providing calibration for variations in ~~ concentrations in recent sediments and ice cores (see Mazaud et aL, 1994). The method developed by Thellier and Thellier (1959) has been widely used as a means of determining paleointensities from igneous rocks. The basis of the method is the assumption that the ratio of the TRM produced in the laboratory to the original TRM (NRM) is equal to the ratio of the laboratory field to the ancient field in which the rock cooled. The method as modified by Coe et al. (1978) and Pr6vot et al. (1985) involves heating the sample to progressively higher temperatures at 10-20~ increments (Fig. 14.10). At each temperature step, the sample is first cooled in field-free space to record the NRM lost; the sample is then reheated at the same temperature and cooled in a known field (typically 40/zT) to record the p-TRM gain. The slope of the line on the so-called Arai plot (Nagata et al., 1963) yields the paleointensity estimate (Fig. 14.10). The bane of the Thellier method is the likelihood of alteration of the magnetic mineralogy during the incremental heating/ cooling cycles. Following Coe et aL (1978), some p-TRM acquisition steps are repeated after heating at higher temperatures (Fig. 14.10), and if the intensity of the p-TRM is reproducible, then it is less likely that alteration has occurred. Other criteria for a reliable paleointensity determination (see Pick and Tauxe, 1993a) include: (1) the plot of pTRM against NRM loss should be linear (Fig. 14.10); (2) susceptibility

14.7 Paleointensity Determinations

25

-

27 |

,~o o

-E 2~ < ~ 15

z

0

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o

9

5

o

-

_

5o9~__a

l

1

I

I

0

5

10

15

p-TRM (x 10.9 Am2)

l

20

/ 14.10 Arai paleointensity plot from the glassy rind of a basaltic pillow basalt (after Pick and Tauxe, 1993a). Figure

should remain constant during the incremental heat treatment; (3) thermal demagnetization of the NRM, measured during the first of the two heating steps, should show a single well-defined component; (4) remanence should be carried by single-domain magnetite. A compilation by Pr6vot et al. (1990) of 12 Mesozoic and Cenozoic Thellier paleointensity determinations (6 of them from papers by Bol'shalov, Solodovnikov, and co-workers) indicates a relative paleointensity low for Early Cretaceous time. Four paleointensity determinations from ODP pillow lavas by Pick and Tauxe (1993b) imply that this paleointensity low continues through most of the Cretaceous with a rather abrupt doubling of the virtual axial dipole moments (VADMs) at the end of the Cretaceous. These results debunk the notion that extended periods of constant geomagnetic polarity (such as the Late Cretaceous long normal) are associated with relatively high paleointensities (e.g., Larson and Olson, 1991). The present paucity of paleointensity data for Cenozoic and Mesozoic time will probably improve considerably in the near future if the recently derived data from volcanic glass (see Pick and Tauxe, 1993a,b) lives up to its early promise. In contrast to the paucity of Thellier paleointensity determinations for pre-Holocene time, there is an abundance of determinations for the last 5 ky with a few determinations dating back to 30 ka (Fig. 14.11). There is considerable scatter in the data due partly to the influence of secular variation, but probably more importantly to alteration of magnetic minerals during the experiment and use of unsuitable archaeological and volcanic materials. Nonetheless, the volcanic data show similarities with the sedimentary relative paleointensity data (Fig. 14.11).

272

14 Rock Magnetic Stratigraphy and Paleointensities

16

~

4

~ ~ ! .: i0

20

i'!

i . .~" ~i~;.'" 30

40

!

i! ~

50

60

70

80

T i m e (ka) Figure 14.| | Comparison of sedimentary relative paleointensity records from Mediterranean Sea (continuous shaded record) with paleointensity data from volcanic rocks and archeomagnetic artefacts (points with error bars) (after Tric et al., 1992).

Deep-sea sediment cores hold considerable promise for giving long continuous records of relative geomagnetic paleointensities. The method relies on the proposition that the intensity of DRM is a linear function of ambient field over the range of geomagnetic field intensity (Kent, 1973; Barton et aL, 1980; Tucker, 1980) and that variations in concentrations of magnetic minerals in the sediment can be compensated by a suitable normalization factor (see review by Tauxe, 1993). The ideal sediment for paleointensity determinations is one in which the remanence is a DRM carried by PSD magnetite in the 1-15/~m grain size range and in which the concentration of magnetite does not vary by more than a factor of 20 or 30 (King et al., 1983). In most studies, the paleointensity record is derived from the NRM intensity (after partial AF demagnetization to eliminate VRM) normalized for variations in magnetite concentration by dividing by ARM, k, SIRM, low-field IRM, or StRM (stirred remanence from redeposition experiments) (Kent and Opdyke, 1977; Tauxe, 1993). More than one normalization factor can be used, strengthening the record in the case of concordance. ARM is often the normalization factor of choice as it is particularly sensitive to SD/PSD magnetite grain sizes, whereas SIRM and k are more sensitive to MD magnetite grains. Meynadier et al. (1992) demonstrated a similarity in coercivity of the NRM and ARM in three piston cores from the Somali Basin, thereby showing that the grain size fraction of magnetite carrying the ARM is similar to that carrying the NRM. For these 5 cm/ky sedimentation rate cores, the NRMz0mT/ARM

14.7 Paleointensity Determinations

273

ratio was used as the measure of relative paleointensity in the 20-140 ka interval. The Somali Basin records are consistent among the three cores and compare well with the 80 ky record of Tric et al. (1992) from the Mediterranean Sea (Fig. 14.11) based on stacked NRM20mT/k data from four cores with sedimentation rates of about 10 cm/ky. These records can be compared with variation seen in Thellier archaeomagnetic and volcanic records for the last 40 ky (e.g., McElhinny and Senanayake, 1982), with relative paleointensities for the last 20 ky from box cores (Constable and Tauxe, 1987), and with longer records from Ontong-Java (Tauxe and Wu, 1990; Tauxe and Shackleton, 1994), Sulu Sea (Schneider, 1993), western equatorial Pacific (Yamazaki and Ioka, 1994), central North Atlantic (Weeks et all., 1995; Lehman et al., 1996), and Labrador Sea (Stoner et al., 1995b). The correlation of the marine paleointensity records to a 84 ky record from Lake Baikal (Peck et al., 1996) is quite straight forward, but correlation to Lau du Bouchet (Thouveny et al., 1990) is less clear. Comparison of the Labrador Sea record with a composite record compiled from the Mediterranean and Somali Basin records (Fig. 14.12) leads to two conclusions. (1) Relative paleointensity records for the last 100 ky can be correlated over large distances, implying that marine sediments can record paleointensities of the geomagnetic dipole field. (2) The AMS 14C age control for the Labrador Sea record is inconsistent with the oxygen isotope age control from the Mediterranean/Somali basins. Paleointensity studies in marine sediments, which hitherto had been restricted to the last ---200 ky, have been greatly expanded by an important study of ODP Leg 138 cores from the equatorial Pacific (Valet and Meynadier, 1993). The record covers the last 3.9 My in cores with mean sedimentation rates not exceeding 2.5 cm/ky (Fig. 14.13). Relative paleointensities were converted into virtual axial dipole moments (VADMs) by calibration to Thellier determinations of paleointensities for the last 0.5 My. Several important features are apparent in the records: (1) paleointensity lows in the Brunhes Chron appear to correlate in age to geomagnetic directional excursions; (2) a sawtooth pattern of paleointensity fluctuations appears to be related to reversal boundaries; reversals tend to coincide with lows in the VADMs with abrupt increases immediately post reversal. The longer the duration of the polarity chron, the greater the abrupt increase in VADM at its onset; and each polarity chron appears to be characterized by decay of VADM leading to the next reversal. This has led to the suggestion that the reversal frequency, and the occurrence of polarity excursions, is ultimately controlled by the intensity of the dipole field, which also controls the stability of a particular polarity state. The correlation of widely spaced sedimentary paleointensity records for the last 140 ky supports the contention that they are reliable records

274

a)

14 Rock Magnetic Stratigraphy and Paleointensities

.~

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Mediterranean Indian Ocean stack 90 I 10

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Individual chronologies (ka)

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Mediterranean Indian Ocean stack 90 ! !0 I

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Labrador Sea chronology (ka)

Figure 14.12 (a) The Labrador Sea relative paleointensity stack (dashed line) with age model derived mainly from AMS 14C dates and a planktic 6180 record. Mediterranean/Indian Ocean (MIO) paleointensity stack (solid line) (Meynadier et aL, 1992; Tric et al., 1992) with ages based on 6180 and tephrochronology. (b) Labrador Sea paleointensity stack correlated to MIO stack using visual ties and plotted on the MIO time scale. (c) MIO paleointensity stack correlated to Labrador Sea stack using visual ties and plotted on the Labrador Sea time scale (after Stoner et al., 1995b).

Figure 14.13 Composite paleointensity (virtual axial dipole moment, VADM) record from ODP Leg 138 (Site 848 and Site 851) in the equatorial Pacific Ocean (after Valet and Meynadier, 1993).

O ,i,,-

00

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276

14 Rock Magnetic Stratigraphy and Paleointensities

2.5

"~ 2 "~ 1.5

Z t~

0.5 Z 0 0

20

40

60

80 100 120 140 160 180 200

Age (ka)

Figure 14.14 Comparison of the relative paleointensity stack (Sint-200, Guyodo and Valet, 1996) with the P-013 paleointensityrecord from the Labrador Sea (Stoner et al., 1995b). The correlation allows the oxygen isotope chronology from Sint-200 to be imported into the Labrador Sea, where obtaining such chronologies directly is problematic.

of the VADM. The presence in the paleointensity record of orbital periodicities (see Meynadier et al., 1992) suggests, however, that climate-controlled lithologic variability may be influencing the records. Tauxe and Wu (1990) advocate use of a coherence function to test whether paleointensity records are coherent (at the 95% confidence level) with the normalizing factor used, and with other magnetic parameters (such as ARM or k) which may be affected by climate-controlled lithologic cycles. Recent paleointensity studies from the Atlantic and Indian Oceans for the Jaramillo to early Brunhes interval (Valet et al., 1994) appear to confirm the general structure of the Leg 138 record (Valet and Meynadier, 1993), indicating that paleointensity records will not only provide us with significant insights into the nature of the geomagnetic field and the mechanism for polarity reversals and polarity excursions but also provide us with a new means of high-resolution stratigraphic correlation. An example of such a correlation is depicted in Figure 14.14. The relative paleointensity stack compiled by Guyodo and Valet (1996), based on paleointensity records that can be directly correlated to the oxygen isotope chronology, is correlated to the paleointensity record from the Labrador Sea (Stoner et al., 1995b) where oxygen isotope data are compromised by lack of benthos and meltwater influence on the planktic record. The correlation in Fig. 14.14 allows the oxygen isotope chronology to be imported into the Labrador Sea, heralding a new era of high resolution magnetic stratigraphy based on geomagnetic paleointensity.

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Index

Advanced piston corer (APC), 99, 126 Afar, 155 African and South American Neogene, 155 Age of Mammals, 160 Alaskan loess, 266 Alpine-Himalayan orogeny, 146 Alternating field demagnetization, 50-51 demagnetizers, 51 Amadeus Basin, 228 Ammonite zones, 186 Ammonitico Rosso facies, 196, 200 Anhysteretic remanent magnetization (ARM), 36, 241,254 Anomalies "x" and "y," 96 Anomaly 32, 168 Ansian/Ladinian boundary, 214 Anthropogenic environmental change, 267 processes, 268 Anticlockwise looping of secular variation, 239 Antiferromagnetism, 28 Anton Chico member, 214 Anza-Borrego, 145 Apparent polar wander path (APWP), 74 Aptian, Albian, and Cenomanian, 125 Aquitanian, 130 Arai plot, 270 Arctic sediments, 160 Arenig and Llanvirn, 228 Argille azzurre, 102 Argon4~ 39 age calibration of the Plio-Pleistocene GPTS, 111 studies, 111

Arikareen and Hemingfordian, 145-146, 156 ARM, 51,252 Astaracian-Vallesian, 146 Astrochronologic calibration, 107 Astrochronological age estimates, 111 Astrochronology, 109 Atlantic margin, 133 Australia, 160, 228, 238 Authigenic hematite, 48 Automated measurement track, 269 Axial dipole moment, 23, 271 geocentric dipole, 15 Bahamas, 101 Bajocian, 200 Baked contact test, 2-5 Bakonycsernye section, 202 Baringo basin, 155 Barremian, 172-173 Barremian/Aptian boundary, 146, 189, 203 Barstovian fossils, 145 Base of the Astaracian, 151 of the Clarendonian, 151 of the PCRS, 220 Basin development, 154 formation and subsidence, 144, 167 Bathonian, 200 Bedding tilt test, 83 Bering land bridge, 163 Berriasian-Aptian magnetic stratigraphy, 191

333

334

Index

Berriasian stratotype section, 193 Berriasian/Valanginian boundary, 203 Big Lost Excursion, 246 Bingham distribution, 60 Biogenic magnetites, 42 Biomagnetostratigraphic correlations, 99 Biostratigraphic control, 191 datums, 105 BIRM, 255 Black Mountain, 228 Blake Excursion, 245 Blancan and Irvingtonian fauna, 145 Blancan LMA, 145 Blocking temperature spectra, 34, 53, 56 Bolucektusi Tepe, 209 Bootstrap, 60, 85 Boreal Realm, 195 Bottom water temperature, 140 Breggia section, 201 Brunhes, 2 Brunhes/Matuyama boundary, 105

CaCO3 cycles, 109 Calabria, 102 Cambrian age, 228 Cambro-Ordovician boundary, 228 Campanian-Maastrictian boundary, 204 Carnian/Ladinian boundary, 212 Carnian/Norian boundary, 208 Cenozoic ~ O record, 136 CENT94, 176 Central Africa, 155 Changes in climate, 267 Changxingian, 216 Chapadmalal, 155 Chattian, 130 Chemical demagnetization, 54 Chemical remanent magnetization

(CRM), 35 Chemostratigraphy, 136 China, 154 Chinese loess, 154, 245 Chron 26n, 132 Chron nomenclature, 78 Chronogram technique, 95 Chronograms, 170 Circle of 95~ confidence (a95), 60 Cismon section, 189

Clarendonian, 145-146 and Barstovian fossils, 145 Clarendonian-Hemphillian, 146 Clarks Fork basin, 159 Climate-controlled lithographic variability, 273, 276 Clockwise looping of secular variation, 239 CM25, 175 Cobb Mountain subchron, 95-96, 249 Cochiti, 96 Coercive force (Hc), 28, 256 Coercivity of remanence, 28, 51 spectra, 56 Coercivity/temperature, 72 Collision of India and Eurasia, 160 Component of magnetization, 54 Concentration of the magnetic mineral, 252 parameter (k), 83 of very fine grained (SP) magnetite, 253 Conglomerate test, 47 Conjugate magnetic anomalies, 117 Continental red beds, 46 Continuous thermal demagnetization of IRM, 70 Core expansion, 251 RC 9-181,109 Correlation of Late Jurassic-Cretaceous Stage, 202 of polarity chrons, 193 with the oxygen isotope curve, 263 Covariance of ~80 values, 140 Cretaceous long normal superchron, 24, 182, 186 magnetic stratigraphy, 182 quiet zone, 230 Cretaceous-Tertiary (K/T) boundary, 119, 159, 182 Critical angle, 83 Cross-correlation technique, 241 Cryogenic magnetometer, 49 Cryptochron, 24, 117 Curie or Ne61 temperature, 28 Curtis formation, 200 Cyclostratigraphy, 107, 125, 186, 207

335

Index

Declination, 9, 16 Deconvolution, 269 Deep Sea Drilling Project (DSDP), 126 Deep sea sediments, 88 Deglaciation of Greenland, 262 Delta 13C record for the Oligocene from DSDP Site 522, 142 record, 140 isotope stages I to IX, 142 180, 130 isotope event Mi7, 164 stratigraphy, 136 time scale, 105 Demagnetization spectra, 55 Deseadan land mammal age, 156 Detrital remanent magnetization (DRM), 35, 45 Deuteric oxidation, 30 Dewey Lake Formation, 218 Diagenetic alteration, 42 growth of maghemite, 43 Diamagnetism, 26 Dinosaur fauna, 159 Dipolar transition fields, 21 Dipole hypothesis, 15, 17 moment, 11 Dolomitization, 101 Donetz basin, 222 Drop in sea level, 162 DSDP Leg 68, 99 72, 130 73, 99 94, 99, 101 Site 609, 130 522, 136, 142 574, 140 Duchesnian, 156 and Chadronian, 156 interchange, 163 Dunkard Formation, 219 subchron, 219

Duration of polarity chrons, 230 Dynamo process, 19

Early and late Barstovian, 151 Aptian, 171-172 Hemingfordian, 164 Early/Middle Triassic, 214 Eocene boundary, 136 Ebro basin, 154 Eccentricity cycles, 109 Elaphas, 164 Ellesmere and Axel Heiberg Islands, 214 End of the PCRS, 219 Environmental changes, 251,268 Eocene/Oligocene boundary, 136, 138, 140 Eolian accumulation rate, 263 detrital magnetite, 44 flux, 264 grain size, 264 Equal duration of ammonite zones, 177 Equus, 156, 164 at Hagermann, 145 Erieu excursion, 243 Eureka Sound Formation, 160 Evidence for reversal of the field, 5 Evolutionary first appearances, 145 Exsolution lamellae, 31

Faunal diachroneity, 160 dispersal, 160 interchange, 160, 166 migration, 144, 164 Fayum, 155 Ferrimagnetic grain-size-sensitive ratios, 260 Ferrimagnetism, 28 Ferromagnetism, 27-28 Field reversal, 3, 4 sampling, 85 Finite rotation poles, 117 First order migratory events, 163 Fisher statistics, 58-60 Flooding of the Mediterranean, 146

336 Formation of unconformities, 162 Frequency-dependent magnetic susceptibility (kf), 253, 255 Future prospects for rock magnetic stratigraphy, 269

Gamma process, 24 Gauss/Matuyama boundary, 155-156, 164 Geomagnetic equator, 12 excursions, 242 field, 14 intensity lows, 23 paleointensity, 276 polarity reversal, 19 polarity time scale (GPTS), 1, 74, 113, 115 pole, 12 Gilsa subchron, 96 Glacial episodes, 140 eustasy, 161,163 Glacial/interglacial variations, 261 Glaciation on Greenland, 142 Global event, 245 Globigerapsis kugleri, 130 Globigerinelloides blowi, 171 Goethite, 33 Gothenberg excursion, 243 Grain size spectrum of ferrimagnetic grains, 263 Grain size, 252 Granulometric parameters, 252 G RAPE, 109 Great Valley Group, 173 Greigite, 34 Griesbachian, 215 Guadalupian or Wordian, 218 Gubbio (Bottacione) section, 182 Gubbio/Contessa sections, 132-133

Hankanina, 136 Hard IRM (HIRM), 252-254, 257, 263 Haute-Savoie, 154 Hauterivian/Barremian boundary, 193 Hawaiian block model, 175 lineations, 169

Index

Heinrich events, 257 Helvetic limestones, 38 Hematite, 31, 62 Hemingfordian-Barstovian boundary, 146 Hemipelagic sediments, 44 Hemphillian, 146, 155 Hemphillian/Blancan boundary, 164 Heterochroneity, 159 Hettangian, 176 Hettangian-Oxfordian, 176 High resolution magnetic stratigraphy, 276 stratigraphic correlation, 276 High-amplitude secular variation, 249 High-frequency field, 253 Hipparion, 146 fauna, 151 HIRM, 253 Hoffman-Day method, 56 Hole 689B, 133 690B, 133 702B, 133 Homo erectus, 155 Huayquerian, 155 Hydraulic piston corer (HPC), 99, 107, 126 Hyperzones, 205 Hysteresis loop, 27, 65

Ice-rafted detritus (IRD), 257 Ice sheet instability, 259 Ice volume change, 136, 140 Illawara reversal, 215 Ilmenite, 30 Immigrant taxa, 161 Inclination, 10 India, 146 Indian Ocean, 130 Induced magnetization, 27 Initial susceptibility, 27 Inner core, 23 Integration of chemostratigraphy and magnetic stratigraphy, 136 Intercontinental faunal exchange, 156, 161 IRM acquisition, 62 ISEA reverse polarity zone, 189 Isoporic charts, 13 Isothermal remanent magnetization (IRM), 36, 254

337

Index

Isotopic stage 19, 105 Isthmus of Panama, 155

Japanese, Hawaiian, and Pheonix Msequence lineations, 168 Jaramillo, 95 creek, 7 Juniata Fro, 225 Jurassic anomalies, 169 magnetic stratigraphy, 196 quiet zone, 169 Jurassic/Cretaceous (Tithonian/Berriasian) boundary, 193

K-Ar ages, 107 K/T Boundary magnetic stratigraphy, 160, 182 Kaena and R6union chrons, 95 k .... 252 karm/k, 255 karmSIRM, 263 Kavur Tepe magnetozones, 208 Kazanian stage, 218 and Tartarian time, 216 Kent and Gradstein timescale, 170 Kiaman magnetic interval, 215 Kimmeridgian-Oxfordian boundary, 175 Kimmeridgian-Tithonian magnetic stratigraphy, 196 Kimmeridgian/Tithonian boundary, 196 Krakow Upland, 200

Lake Baikal, 273 sediment (magnetic properties), 45, 265 Tahoe, 244 Windermere, 233 Land mammal ages (LMAs), 145 Laschamp excursion, 244 Late Arikareen, 164 Eocene, 156 Miocene carbon shift, 142, 164 Least squares fitting, 56 Leg 142, 152

Lena River, 228 Lengths of magnetic intervals, 230 Lineated magnetic anomalies, 113, 168 Llanvirn-Caradoc, 225 Lock in zone, 39 Loess deposits, 46, 265 Low-frequency field, 253 Lowrie-Fuller test, 66

M0-M29, 169 M-sequence magnetic anomalies, 81,191 polarity chrons, 169, 174, 191 Maghemite, 33 Maghemitization, 31 Magnetic anomalies, 111, 115 enhancement, 265, 267 epochs, 5, 78 equator, 10 excursions, 20, 242 grain size (from ARM/k), 264 mineral flux, 269 mineralogy, 253 minerals, 28-29 parameters, 251 polarity stratigraphy, 1 pole, 10 properties of marine sediments, 37 stratigraphy, 75 susceptibility (k), 250-251 anisotropy, 25 Magnetite, 29 Magnetization components, 50 Magnetobiochemochronology, 143 Magnetobiochronology, 81, 132, 136, 143 Magnetometers, 4 Magnetostratigraphic classification, 75 polarity classification, 75 reversal horizons, 75 units, 75, 78 section, 119 studies of terrestrial sequences, 144 type section, 81,205 units, 75 Magnetozone, 76

338 Maiolica Limestone Formation, 191 Mammal dispersal, 160 faunas, 159, 161 Mammalian biochronology, 145 Mammoth subchron, 95 Mammuthus in North America, 145 Mantle plumes, 232 Marine records of secular variation, 238 sediments, 253 Mariposa Formation, 175 Massif Central, France, 244 Massignano section, 136 Master secular variation curve for Britain, 234, 239 Matuyama, 3 Mauch Chunk Formation, 222 Maud Rise (ODP Leg 113), 133 Maximum angular deviation (MAD), 56 Mayerling section (Austria), 212 McLaughlin cycles, 207 Melt event, 261 Mesozoic platform limestone, 38 Mil, 140, 161 Mi7, 140, 161 Mid-Mesozoic M-sequence, 168 Mid-Miocene 8180 shift, 140 Middle Albian polarity zones, 189 Oligocene, 164 Silurian, 225 Middle/late Eocene, 136 Midland Valley of Scotland, 222 Migration of hydrocarbons, 38 Migrationary first appearances, 145, 163 Migratory events, 163-164 Milankovitch cycles, 140 cyclostratigraphy, 207, 208 Miocene magnetic stratigraphy, 126 magnetobiochronology, 130 Miocene-Pliocene boundary, 102, 105, 146 MN1, 154 Moenkopi Formation, 212 Mono Lake excursion, 234, 243 Montehermosan, 155 Monte Narbone Formation, 102 Monterey event, 142 Morin transition, 70

Index

Morrison Formation, 200 Morrowan paleosols, 220 MP29, 154 MP30, 154 MP30/MN1 boundary, 154 Mu metal shields, 54 Mururoa Atoll, 101 Muschelkalk Formation of Spain, 214

Namurian, 220 Nannoconid crisis, 172 Nannofloral datums, 202 Nannofossil biostratigraphy, 132, 186 Narabeen Chocolate Shales, 215 Neogene land mammal ages, 161 Newark Basin, 207 Nondipole components, 239 Nonmarine sediments, 144 Normalization factor, 272 North American and Eurasian Paleogene, 156 neogene and quaternary, 145 Northern Apennine foredeep, 105 Hemisphere Glaciation, 256, 265 NRM, 49, 254 analysis, 61 NRM/ARM, 241 Numerical age control, 119, 171 Nunivak subchron, 96

Ocean Drilling Program (ODP), 126 Ochoan, 218 Octupole, 24 transition field geometrics, 19 ODP Holes 845A and 845B, 130 Leg 107, 102, 138, 109 Site 710, 130 Oil and Oi2, 140, 161 Olduvai chron, 95 Gorge, 7 Oligocene, 156 Oligocene-Miocene boundary, 130, 154 Omo area, 155

339

Index

Orbital cycles of eccentricity, obliquity, and precession. 107 Ordovician, 225 Orientation, 49 Orleanian/Astaracian boundary, 152 Orthogonal projections, 54 Oxfordian, 176 Oxfordian-Aptian time scales, 175 Oxfordian-Callovian magnetic stratigraphy, 200 Oxfordian/Kimmeridgian boundary, 196, 203 Oxygen isotope zones, 140, 161

P-wave propagation, 23 Pacific anomalies, 168 Pakistan, 146 Paleocene-Eocene boundary, 136 Paleoclimatology, 160 Paleogene age, 159 magnetic stratigraphy, 132 Paleogeography, 160 Paleointensity determinations, 270 high, 241 lows, 273 magnetic stratigraphy, 2, 273 Paleomagnetic pole, 16 Paleoprecipitation, 265 Paleogene magnetic stratigraphy, 132 Paramagnetic ions, 26 Paramagnetism, 26 Parameter estimation, 84 Pass through cryogenic magnetometers, 87-88, 234 Passive margin unconformities, 161 Paterson Volcanics, 220 Pattern fit, 91 Pedogenic and detrital susceptibility, 265 Pelagic deep-sea sediments, 101 limestones, 39 Periodicity of ---100 ky, 105 Permo-Carboniferous Reverse Superchron (PCRS), 215 Permo-Triassic boundary, 214-215 Physical disturbance of sediments, 243 Pictou sediments, 219

Piston core collections, 99 Platform carbonates, 101 Pliensbachian-early Bajocian interval, 201 Plio-Pleistocene boundary, 102 climatic cycles, 154 glaciation, 145 lake and loess sequences, 145 land mammal sequences, 144 magnetic stratigraphies, 101 polarity chrons, 95 Polarity, 9 bias, 215, 229, 232 chron nomenclature, 75 transition records and VGP paths, 20 zone, 76 Potwar basin, 167 Pre-Callovian Jurassic magnetic stratigraphy, 200 Precessional forcing, 125 variance, 263 Precision parameter, 59 Preferred longitudinal VGP paths, 21 Pringle Falls excursion, 245 Probability density P, 58 Progressive deforestation, 267 Pyrrhotite, 34

Quadrupole, 24 Quality criteria, 93 Quiet zone, 168

Rates of sediment accumulation, 88 Red clay facies, 43 Relative geomagnetic paleointensities, 272 Relaxation time, 34, 51 Remagnetization circles, 56 Remanent coercivity, 28 magnetization, 27 Repetitive waveform, 235 Reproducibility of magnetozones, 89 Resolution of astrochronological timescale, 112 Resultant direction (R), 59 R6union event, 96

340

Reversal frequency, 208 process, 249 test, 82 Reversals per million years, 232 Ringerike Formation, 223 Rock magnetic stratigraphy, 2, 250 in lake sediments, 266 Rossello composite, 102 Rotliegende and Zechstein Formations, 218 Ruscinian, 146

S ratios, 252, 255 Salla beds of Bolivia, 156 Salt Range, 216 Sample internal, 87 Sampling unindurated sediments, 89 San Juan Basin of New Mexico, 159 San Pedro Valley, 144 Santa Catalina Basin, 238 Santonian-Campanian, 203 Santonian-Maastrichtian magnetic stratigraphy, 186 Sapropels, 109 Saturated bulk (GRAPE) density, 250-251 Saturation isothermal remanence (SIRM), 241,252 magnetization (Ms), 28, 62, 255 remanent magnetization, 28 Sawtooth pattern, 273 Sea floor spreading, 113 Sea level fall, 161, 164 Second-order migratory events, 163-164 Secondary magnetizations, 232 Secular variation, 13-14, 233 magnetic stratigraphy, 2, 233 Sediment records of secular variation, 234 Sedimentary completeness, 167 environments, 263 processes, 154, 167 relative paleointensity data, 271 Sedimentology, 144 Self-reversal, 4 mechanism, 37 Separation techniques, 72 Sequence boundaries, 164 stratigraphy, 162-163

Index

Shallow water carbonates, 37 Shungura sequence, 155 Siberian Platform, 225, 228 Sidufjall, 96 SIRM, 252 SIRM/k, 255 SIRM/k . . . . 255 Site 607, 108 677, 107-108 846, 107 Siwalik sediments of Pakistan and India, 47, 146, 154 Slump test, 85 Smoothing of individual records, 242 Solar dipole field, 20 South America, 160, 166 South Atlantic, 132 DSDP Sites 522, 523, and 524, 132 Southern England, 133 Spain, 146 Spectral analysis of CaCO3, 109 Spinner magnetometer, 50 Spline functions, 242 Spurious data, 243 SQUID, 49 Stage 1,105 Statistics, 58 Stirred remanent magnetization (StRM), 241 Stratigraphic completeness, 167 Stratotype sections, 102 Strontium87/Strontium86, 142 stratigraphy, 130 Subchrons, 95 Summer Lake, 244 Summerville Formation, 200 Susceptibility, 26 variations in the Chinese loess, 265 Susceptibility/temperature, 71 Synfolding magnetization, 83

Tabianian and Piacenzian, 102 Tectonic development, 166 events, 166 history, 144 Tectonics, 154

Index

Tephrachronology and 61sO stratigraphy, 245 Terrestrial magnetic stratigraphy, 144 sediments, 88 Terrigeneous (eolian) flux, 256 Tethyan sections, 196 TG20 and TG22, 164 Thellier and Thellier method, 270 Thermal demagnetization, 51 remanent magnetization (TRM), 4, 34 Three-axis IRM, 65 Thulean land bridge, 163 Thvera, 96 Tiffanian, 159 Time averaged dipole, 15 versus depth plot, 248 Timing remanence acquisition, 82 Tiny wiggles, 24, 117 Titanohematite, 31, 33 Titanomaghemite, 33 Titanomagnetite, 29, 61 Tithonian/Berriasian boundary, 203 Toadstool Park-Roundtop section in Nebraska, 159 Toarcian stratotype, 202 Tommotian-Atdabanian, 228 Torrejonian faunas, 159 Triassic magnetic stratigraphy, 207 Trubi Limestones, 102 Marls, 109 Turbiditic "Flysch" deposits, 44 Turkey, 146 Turolian-Ruscinian boundary, 146 Tyrrhenian Sea, 102

U-channels, 88, 269 Unblocking temperature, 52

34 ]

Uplift episodes, 167 Upper Normal Interval, 207 Permian type section, 218 Uquian land mammal age, 155

Valanginian and Berriasian, 175 Valanginian/Hauterivian boundary, 193 Vallesian/Turolian boundary, 146 Vastergotland, 225 Vector loops, 240 Vertebrate evolution and migration, 154 paleontology, 144 Verwey transition, 69 VGP latitude, 91 Vieja section of west Texas, 156 Virtual geomagnetic pole, 16, 21 Viscous remanent magnetization (VRM), 36 Visean, 220, 222 Volcanic air fall ashes, 132 Volumetric magnetic susceptibility, 254 Vrica, Italy, 102, 105

Wargal and Chidru Formations, 216 Wasp-waisted hysteresis, 39 Watson test for randomness, 85 Westward drift, 13, 237, 239 rate, 240 Whitneyan/Arikareen boundary, 159

Zanclean, 102 Zijderveld diagram, 54-56

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International Geophysics Series E D I T E D BY

RENATA DMOWSKA Division of Applied Science Harvard University Cambridge, Massachusetts JAMES R. H O L T O N Department of Atmospheric Sciences University of Washington Seattle, Washington

Volume 1

BENOGUTENBERG. Physics of the Earth's Interior. 1959"

Volume 2

JOSEPHW. CHAMBERLAIN.Physics of the Aurora and Airglow. 1961"

Volume 3

S.K. RUNCORN (ed.) Continental Drift. 1962"

Volume 4

C. E. JUNGE. Air Chemistry and Radioactivity. 1963"

Volume 5

ROBERT G. FLEAGLE AND JOOST A. BUSINGER. A n Introduction to Atmospheric

Physics. 1963" Volume 6

L. DUFOUR AND R. DEFAY. Thermodynamics of Clouds. 1963"

Volume 7

H . U . ROLL. Physics of the Marine Atmosphere. 1965"

Volume 8

RICHARDA. CRAIG. The Upper Atmosphere: Meteorology and Physics. 1965"

Volume 9

WILLISL. WEBB. Structure of the Stratosphere and Mesosphere. 1966"

Volume 10

MICHELE CAPUTO. The Gravity Field of the Earth from Classical and Modern Methods. 1967"

Volume 11

S. MATSUSHITAAND WALLACEH. CAMPBELL(eds.) Physics of Geomagnetic Phenomena. (In two volumes.) 1967"

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K. YA KONDRATYEV. Radiation in the Atmosphere. 1969"

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Volume 13

E. PALI~IENAND C. W. NEWTON. Atmospheric Circulation Systems: Their Structure and Physical Interpretation. 1969"

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HENRYRISHBETHAND OWES K. GARRIO'IT.Introduction to Ionospheric Physics. 1969"

Volume 15

C.S. RAMAGE. Monsoon Meteorology. 1971"

Volume 16

JAMESR. HOLTON. An Introduction to Dynamic Meteorology. 1972"

Volume 17

K.C. YEn AND C. H. LIU. Theory of Ionospheric Waves. 1972"

Volume 18

M . I . BUDVKO. Climate and Life. 1974"

Volume 19

MELVINE. STERN. Ocean Circulation Physics. 1975

Volume 20

J . A . JACOBS. The Earth's Core. 1975"

Volume 21

DAVIDH. MILLER.Water at the Surface of the Earth: An Introduction to Ecosystem Hydrodynamics. 1977

Volume 22

JOSEPHW. CHAMBERLAIN.Theory of Planetary Atmospheres: An Introduction to Their Physics and Chemistry. 1978"

Volume 23

JAMESR. HOLTON. An Introduction to Dynamic Meteorology, Second Edition. 1979"

Volume 24

ARNETYS. DENNIS. Weather Modification by Cloud Seeding. 1980

Volume 25

ROBERTG. FLEAGLEAND JOOST A. BUSINGER. An Introduction to Atmospheric Physics, Second Edition. 1980

Volume 26

KuG-NAN LIOU. An Introduction to Atmospheric Radiation. 1980

Volume 27

DAVID H. MILLER. Energy at the Surface of the Earth: An Introduction to the Energetics of Ecosystems. 1981

Volume 28

HELMUTG. LANDSBERG.The Urban Climate. 1981

Volume 29

M . I . BUDYKO. The Earth's Climate: Past and Future. 1982"

Volume 30

ADRIANE. GILL. Atmosphere-Ocean Dynamics. 1982

Volume 31

PAOLOLANZANO. Deformations of an Elastic Earth. 1982"

Volume 32

RONALDZ. MERRILLAND MICHAELW. MCELHINNY.The Earth's Magnetic Field. Its History, Origin, and Planetary Perspective. 1983

Volume 33

JOHN S. LEWIS AND RONALD G. PRINN. Planets and Their Atmospheres: Origin and Evolution. 1983

Volume 34

ROLE MEISSNER. The Continental Crust: A Geophysical Approach. 1986

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M . U . SAGITOV,B. BODKI, V. S. NAZARENKO, AND KH. G. TADZHIDINOV. Lunar Gravimetry. 1986

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Volume 36

JOSEPHW. CHAMBERLAINAND DONALD M. HUNTEN. Theory of Planetary Atmospheres: An Introduction to Their Physics and Chemistry, Second Edition. 1987

Volume 37

J . A . JACOBS. The Earth's Core, Second Edition. 1987

Volume 38

J . R . APEL. Principles of Ocean Physics. 1987

Volume 39

MARTINA. UMAN. The Lightning Discharge. 1987

Volume 40

DAVIDG. ANDREWS, JAMES R. HOLTON, AND CONWAY B. LEOVY. Middle Atmosphere Dynamics. 1987

Volume 41

PETERWARNECK. Chemistry of the Natural Atmosphere. 1988

Volume 42

S. PAL ARYA. Introduction to Micrometeorology. 1988

Volume 43

MICHAELC. KELLEY. The Earth's Ionosphere. 1989

Volume 44

WILLIAMR. COTTONAND RICHARDA. ANTHES. Storm and Cloud Dynamics. 1989

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WILLIAMMENKE. Geophysical Data Analysis: Discrete Inverse Theory, Revised Edition. 1989

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S. GEORGE PHILANDER. E1 Nifio, La Nifia, and the Southern Oscillation. 1990

Volume 47

ROBERTA. BROWN. Fluid Mechanics of the Atmosphere. 1991

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JAMES R. HOLTON. An Introduction to Dynamic Meteorology, Third Edition. 1992

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ALEXANDERA. KAUFMAN. Geophysical Field Theory and Method. Part A: Gravitational, Electric, and Magnetic Fields, 1992 Part B: Electromagnetic Fields I. 1994 Part C: Electromagnetic Fields II. 1994

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SAMUELS. BUTCHER, GORDON H. ORIANS, ROBERT J. CHARLSON, AND GORDON V. WOLFE. Global Biogeochemical Cycles. 1992

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BRIANEVANSAND TENG-FONGWONG. Fault Mechanics and Transport Properties of Rocks. 1992

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ROBERTE. HUFFMAN. Atmospheric Ultraviolet Remote Sensing. 1992

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ROBERTA. HOUZE, JR. Cloud Dynamics. 1993

Volume 54

PETERV. HOBBS. Aerosol-Cloud-Climate Interactions. 1993

Volume 55

S.J. GIBoWICZ AND A. KIJKO. An Introduction to Mining Seismology. 1993

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DENNIS L. HARTMANN. Global Physical Climatology. 1994

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MICHAELP. RYAN. Magmatic Systems. 1994

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THORNELAY AND TERRY C. WALLACE. Modern Global Seismology. 1995

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DANIELS. WILKS. Statistical Methods in the Atmospheric Sciences. 1995

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FREDERIKNEBEKER. Calculating the Weather. 1995

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MURRVL. SALBV.Fundamentals of Atmospheric Physics. 1996.

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JAMESP. MCCALPIN. Paleoseismology. 1996.

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RONALDT. MERRILL,MICHAELW. MCELHINNY,AND PHILLIPL. MCFADDEN. The Magnetic Field of the Earth: Paleomagnetism, the Core, and the Deep Mantle. 1996.

ISBN 0-12--527470-X 90040

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